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Digitized by the Internet Archive 
in 2014 

Carf/L tlh oul 









of Nature 



de la nature 

Above and cover: 

Medallion struck in 
southern Germany 

in memory of the 

great famine of 1816-1817. 

The inscription reads: 
"Great is the distress, 
Oh Lord, have pity." 

Both faces shown; from Volcano Weather, 
The Story of 1816, The Year Without a Summer 
by Henry and Elizabeth Stommel 




OTTAWA, 1992 

MAY 9 1 1992 


©1992 Canadian Museum of Nature ©1992 Musée canadien de la nature 
Published by the: Publié par le : 

Canadian Museum of Nature Musée canadien de la nature 

Ottawa, Canada K1P 6P4 Ottawa, Canada K1P 6P4 

Catalogue No. NM95-20/1 1991-E N° de catalogue NM95-20/1 1991-E 
Available by mail order from: L’éditeur remplet les commandes postales 

adressées au : 

Canadian Museum of Nature Musée canadien de la nature 
Direct Mail Section Section des commandes postales 
P.O. Box 3443, Station "D" C.P. 3443, succursale D 
Ottawa, Canada K1P 6P4 Ottawa, Canada K1P 6P4 
Printed in Canada Imprimé au Canada 
ISBN: 0-660-13063-7 ISBN : 0-660-13063-7 

Text pages printed on paper Les pages du texte sont imprimés 

containing recycled fibre. sur un papier contenant 

des fibres recyclés. 

Print of original handwritten copy of Lord Byron’s poem "Darkness" 


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Kit men fog ho fifeim CE Bask 
[ & Be Li, lkin — ral el heigl 

Wan LMS gel’ ae Maar a 
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. Written 
at Geneva during 1816 (courtesy of Princeton University Library). 


I HAD a dream, which was not all a dream. 

The bright sun was extinguish’d, and the stars 
Did wander darkling in the eternal space, 
Rayless, and pathless, and the icy earth 

Swung blind and blackening in the moonless air ; 
Morn came and went—and came, and brought no day, 
And men forgot their passions in the dread 

Of this their desolation ; and all hearts 

Were chill’d into a selfish prayer for light : 

And they did live by watchfires—and the thrones, 
The palaces of crowned kings—the huts, 

The habitations of all things which dwell, 

Were burnt for beacons ; cities were consumed, 
And men were gather’d round their blazing homes 

To look once more into each other's face ; 

Printed version of Lord Byron’s poem "Darkness" (courtesy of Princeton University Library). 



C.R. Harington 

Before Tambora: the Sun and Climate, 1790-1830 
John A, Eddy 

Eyewitness Account of the Distant Effects of the Tambora Eruption 
of April 1815 
Michael R. Rampino 

The Eruption of Tambora in 1815: Environmental Effects and 
Eruption Dynamics 
Haraldur Sigurdsson and Steven Carey 

The Possible Effects of the Tambora Eruption in 1815 on Atmospheric 
Thermal and Chemical Structure and Surface Climate 
R.K.R. Vupputuri 

Climatic Effects of the 1783 Laki Eruption 
Charles A. Wood 

The Effects of Major Volcanic Eruptions on Canadian Surface Temperatures 
Walter R. Skinner 

Northern Hemisphere 

North America 
Climate in 1816 and 1811-20 as Reconstructed from Western North American 
Tree-Ring Chronologies 
J.M. Lough 

Volcanic Effects on Colorado Plateau Douglas-Fir Tree Rings 
Malcolm K. Cleaveland 

1816 in Perspective: the View from the Northeastern United States 
William A. Baron 









Expansion of Toronto Temperature Time-Series from 1840 to 1778 Using 
Various United States and Other Data 145 
R.B. Crowe 

Climate in Canada, 1809-20: Three Approaches to the Hudson’s Bay 
Company Archives as an Historical Database 162 
Cynthia Wilson 

Climatic Change, Droughts and Their Social Impact: Central Canada, 
1811-20, a Classic Example 185 
Timothy F. Ball 

The Year without a Summer: Its Impact on the Fur Trade and History 
of Western Canada 196 
Timothy F. Ball 

The Ecology of a Famine: Northwestern Ontario in 1815-17 203 
Roger Suffling and Ron Fritz 

The Development and Testing of a Methodology for Extracting Sea-Ice 
Data from Ships’ Log-Books 218 
Marcia Faurer 

River Ice and Sea Ice in the Hudson Bay Region during the Second 
Decade of the Nineteenth Century 238 
A.J.W. Catchpole 

The Climate of the Labrador Sea in the Spring and Summer of 1816, 
and Comparisons with Modern Analogues 245 
John P. Newell 

Spatial Patterns of Tree-Growth Anomalies from the North American 
Boreal Treeline in the Early 1800s, Including the Year 1816 255 
Gordon C. Jacoby, Jr. and Rosanne D’Arrigo 

Early Nineteenth-Century Tree-Ring Series from Treeline Sites 
in the Middle Canadian Rockies 266 
B.H. Luckman and M.E. Colenutt 

How Did Treeline White Spruce at Churchill, Manitoba. Respond 
to Conditions around 1816? 281 
David C. Fayle, Catherine V. Bentley and Peter A. Scott 

The Climate of Central Canada and Southwestern Europe Reconstructed 
by Combining Various Types of Proxy Data: a Detailed Analysis of 
the 1810-20 Period* 291 
J. Guiot 

Climatic Conditions for the Period Surrounding the Tambora Signal 
in Ice Cores from the Canadian High Arctic Islands 
Bea Taylor Alt, David A. Fisher and Roy M. Koerner 

Europe (including Iceland) 
1816 - a Year without a Summer in Iceland? 
A.E.J. Ogilvie 

First Essay at Reconstructing the General Atmospheric Circulation 
in 1816 and the Early Nineteenth Century 
H.H. Lamb 

Weather Patterns over Europe in 1816 
John Kington 

The Climate of Europe during the 1810s with Special Reference to 1816 
K.R. Briffa and P.D. Jones 

The 1810s in the Baltic Region, 1816 in Particular: Air Temperatures, 
Grain Supply and Mortality 
J. Neumann 

The Years without a Summer in Switzerland: 1628 and 1816 
Christian Pfister 

Climatic Conditions of 1815 and 1816 from Tree-Ring Analysis in the 
Tatra Mountains 
Zdzistaw Bednarz and Janina Trepinska 

Major Volcanic Eruptions in the Nineteenth and Twentieth Centuries 
and Temperatures in Central Europe 
Vladimir Briizek 

Climate over India during the First Quarter of the Nineteenth Century 
G.B. Pant, B. Parthasarathy and N.A. Sontakke 

Evidence for Anomalous Cold Weather in China 1815-17 
Pei-Yuan Zhang, Wei-Chyung Wang and Sultan Hameed 

Was There a Colder Summer in China in 1816? 
Huang Jiayou 

The Reconstructed Position of the Polar Frontal Zone around Japan 
in the Summer of 1816 
Yasufumi Tsukamura 

The Climate of Japan in 1816 as Compared with an Extremely Cool 
Summer Climate in 1783 
T. Mikami and Y. Tsukamura 














Southern Hemisphere 

Evidence for Changes in Climate and Environment in 1816 as 
Recorded in Ice Cores from the Quelccaya Ice Cap, Peru, the 
Dunde Ice Cap, China and Siple Station, Antarctica* 

Lonnie G. Thompson and Ellen Mosley-Thompson 

Changes in Southern South American Tree-Ring Chronologies 
following Major Volcanic Eruptions between 1750 and 1970 
Ricardo Villalba and Jose A. Boninsegna 

Tree-Ring Chronologies from Endemic Australian and New Zealand 
Conifers 1800-30 
Jonathan Palmer and John Ogden 

New Zealand Temperatures, 1800-30 
David A. Norton 


Workshop on World Climate in 1816: a Summary and Discussion of 
Cynthia Wilson 










* The geographic sections above are not exact. For example, J. Guiot’s paper, although listed 
under North America, also provides substantial information on southwestern Europe and 
northwestern Africa. Similarly, the paper by L. Thompson and E. Mosley-Thompson, although 

listed under Southern Hemisphere, also concerns China. 


The editor is grateful to his colleagues on the Organizing Committee of the international meeting 
("The Year Without a Summer? Climate in 1816", Ottawa, 25-28 June 1988) from which this 
volume arose: Drs. C. Wilson, A.J.W. Catchpole, T.F. Ball, R.M. Koerner, G.C. Jacoby 
(Members); Mrs. Gail Rice (Secretary-Treasurer); and Mr. Kieran Shepherd (Coordinator, Poster 
Presentations). I am also grateful to the following institutions for so firmly supporting the 
meeting: Canadian Climate Centre; Climatic Research Unit, University of East Anglia; National 
Center for Atmospheric Research (operated by the University Corporation for Atmospheric 
Research under the sponsorship of the United States National Science Foundation); and the World 
Meteorological Organization. I thank the directorate of the museum for its interest in and 
encouragement of the project. 

Joanne Dinn (Paleobiology Division) and Marie-Anne Resiga helped greatly in preparing this 
book, as did Mireille Boisonneau, Arch Stewart (Canadian Museum of Nature Library) and 
Daphne Sanderson (Canadian Climate Centre Library). Sharon Helman kindly redrafted several 
of the figures, and with Bonnie Livingstone (Publications Division) provided strong support 
during the last phases of preparing this volume. 

Finally, I express my sincere thanks to Cynthia Wilson and Tim Ball for their help in organizing 
the Workshop, as well as to Richard Martin for audiotaping the discussions. Cynthia Wilson 
performed a particularly useful service by analyzing and summarizing the Workshop results. 


This book is the last gasp of the National Museum of Natural Sciences (now Canadian Museum 
of Nature) Climatic Change in Canada Project! Because of Canada’s vulnerability to climatic 
change, and the lack of an integrated multidisciplinary program for studying our past climate this 
project was organized. Since its beginning in 1977, a basic aim of the project has been to publish 
in our Syllogeus series significant data on climatic change in Canada since the peak of the last 
glaciation (about 20,000 years ago). 

We began the project with a general assessment of Quaternary paleoclimatic information available 
in Canada and techniques that could be used for interpreting it (Syllogeus 26, 1980); we later 
broadened the number of disciplines involved and actually began gathering and interpreting the 
paleoclimatic data (Syllogeus 33, 1981; 49, 1983); and then produced an annotated bibliography 
on the subject (Syllogeus 51, 1984). In May 1983, the project sponsored an international meeting 
"Critical Periods in the Quaternary Climatic History of Northern North America" (Syllogeus 55, 
1985). It was clear from papers in Syllogeus 55 that several authors had gone well beyond the 
data-gathering stage: Alan Catchpole not only tested the value of one type of proxy data (climatic 
records from Hudson’s Bay Company documents, including Ships’ logs) against another (Marion 
Parker’s tree-ring records) for the Hudson Bay region, but showed that sea-ice conditions were 
indicative of prevailing northerly or northwesterly winds, pumping cold arctic air over the central 
and eastern parts of North America in the summer of 1816; and Cynthia Wilson took a 
magnificent step forward by providing a series of six daily weather maps for early June 1816 - 
actually showing the tracks of high- and low-pressure areas across central and eastern North 

These papers prompted me to consider convening an international meeting focusing on global 
climate during 1816, "the year without a summer". What were conditions like beyond the regions 
so well documented by John D. Post in The Last Subsistence Crisis in the Western World (Johns 
Hopkins University Press, Baltimore, 1977) and Henry and Elizabeth Stommel in Volcano 
Weather, the Story of 1816, the Year without a Summer (Seven Seas Press, Newport, 1983), 
among others? Could anything useful on a global basis be added to Lamb’s and Johnson’s (1966, 
Figure 5) excellently constructed pressure map for July 1816 extending from western Europe 
across the Atlantic Ocean to central North America (Lamb, this volume)? 

Accordingly, I wrote to Professor Hubert Lamb in Norwich in February 1985 and received an 
encouraging, constructive reply: "Your idea of holding some sort of a conference or workshop 
meeting specifically to put together the best possible reconstruction of summer 1816, or of the 
whole ‘year without a summer’, or perhaps usefully rather more of that decade particularly aimed 
at covering the period from just before the atmospheric/radiation budget disturbance caused by 
the huge volcanic eruption of Tambora in 1815 till the return to the status quo ante, has intriguing 
possibilities." The proposal for this meeting was approved by the director of the museum in 1986. 

The objective of the meeting was, by bringing together workers in various fields (e.g., 
volcanologists, glaciologists, climatologists, tree-ring experts, geographers, historians and 
biologists) from various countries, to gain the clearest picture possible of weather and climatic 
sequences in different parts of the world during 1816, or about that time (e.g., 1810-20), in an 
effort to discover key factors influencing the unusual weather then. For example, how important 

was the eruption of Tambora, and what other cooling influences may have been involved? How 
widespread were the cold summer conditions from a global viewpoint? Did blocking play an 
important part? 

From the beginning, the Workshop was considered to be the heart of the meeting. The attempt 
to actually plot weather and climatic data from various sources for the Tambora period on base 
maps proved challenging, frustrating and exciting. Could we really shed more light on the nature 
of the climatic events, their intensity and timing? Although evidence is circumstantial, it seems 
that widespread cooling was underway before the eruption of Tambora. Evidently, the massive 
injection of Tambora aerosols into the atmosphere in 1815 resulted in crossing a threshold to 
highly anomalous weather (probably involving blocking highs and break monsoons) in many parts 
of the globe. Certainly "the year without a summer" in 1816 was a regional phenomenon. In the 
northern hemisphere parts of western North America, eastern Europe and Japan seem to have had 
average or above-average temperatures, as opposed to the remarkable cold that characterized 
much of eastern North America, western Europe, and China. Incursion of freezing arctic air 
southward in one region was offset by poleward flow of tropical air in another. In the southern 
hemisphere, El Nifio may have diminished the cooling reflected in tree-ring records from 
Argentina in 1816-17, whereas in 1817-18 the tremendous moderating influence of the Pacific 
Ocean may have effectively damped any cooling recorded there (see Wilson, Workshop section, 
for more details on the group’s findings). 

This book is intended for those who are deeply interested in: historical climate (particularly that 
of the Little Ice Age) and its human impact; relationships between volcanism and climate; and 
the ways paleoclimatic proxy data are gathered, treated and interpreted. The volume begins with 
a general section concerning: solar influences on the trend of climate before the eruption of 
Tambora; a vivid eyewitness account of the eruption; the nature of the eruption, the aerosol 
produced and its course through the atmosphere - as well as a discussion of the effects of the 
1783 eruption of Laki in Iceland on climate for comparative purposes and a consideration of the 
effects of major volcanic eruptions following Krakatau (1883) on Canadian temperatures. 
Coverage is then (loosely) geographic, first dealing with the northern hemisphere (North 
America, Europe, Asia), then the southern hemisphere (South America, Antarctica, Australia and 
New Zealand). Perhaps readers will gather from these contributions an inkling of the tremendous 
investment in time that is presently required to distil a useful drop of paleoclimatic data from 
archival and other sources. 

Finally, I hope that this exercise will lead others to look more carefully at the "Tambora period" 
and similar paleoclimatic problems - adding data in vast expanses of the globe where our evidence 
is deficient, as well as testing and refining data given here until a more coherent picture emerges. 
Information presented in this volume may also be food for ravenous paleoclimatic modellers! 

C.R. Harington 

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Before Tambora: The Sun and Climate, 1790-1830 

John A. Eddy’ 


The unusual summer of 1816 is commonly attributed to the increase in atmospheric turbidity that 
followed the eruption of Mount Tambora (Stommel and Stommel 1979). The awesome eruption 
occurred, in fact, during a span of several decades of colder climate that had interrupted the 
gradual global warming that followed seventeenth century extrema of the Little Ice Age (Lamb 
1985). These background trends may well explain a particularly severe seasonal response in 1816 
to a short-term injection of volcanic dust. The colder climate that characterized the opening 
decades of the nineteenth century was quite possibly related to a coincident depression in solar 
activity between about 1790 and 1830, called the "Dalton Minimum" or sometimes the “Little 
Maunder Minimum" (Siscoe 1980). The probability of a solar connection is strengthened by 
recent analyses of long-term changes in the level of solar activity and decadal averages of global- 
surface temperature in the last 100 years (Reid and Gage 1987), as well as in the correspondence 
of the Maunder Minimum in solar activity (1645-1715). A probable mechanism for solar forcing 
can be found in recent spaceborne measurements of year-to-year variations in the so-called "solar 
constant" (Willson et al. 1986). I plan to examine the evidence for solar and climatic anomalies 
in the period from about 1790-1830 and the recent findings that provide a probable connection 
between the sun and long-term climatic change. 

Eddy, J.A. 1977. The case of the missing sunspots. Scientific American 236:80-92. 

Lamb, H.H. 1985. Climate History and the Future. Princeton University Press, Princeton, 
New Jersey. 884 pp. 

Reid, G. and K.S. Gage. 1987. Influence of solar variability on global sea surface 
temperatures. Nature 329(6135):142-143. 

Siscoe, G.L. 1980. Evidence in the auroral record for secular solar variability. Review of 
Geophysics and Space Physics 18:647-658. 

Stommel, H. and E. Stommel. 1979. The year without a summer. Scientific American 
240: 176-186. 

Willson, R.C., H.S. Hudson, C. Frohlich and R.W. Brusa. 1986. Long-term downward trend 
in total solar irradiance. Science 234:1114-1117. 

' University Corporation for Atmospheric Research, Boulder, Colorado 80307, U.S.A. 


Eyewitness Account of the Distant Effects of the Tambora Eruption of 
April 1815 

Michael R. Rampino’ 


The following is a brief description of the effects of the eruption of Tambora volcano in 1815 on 
conditions about 800 km away in eastern Java. Evidently Tambora was quite active for at least 
six days prior to the cataclysmic eruption of 11 April 1815, and direct cooling was associated — 
with the ash cloud. 


Large explosive volcanic eruptions can have far-reaching effects on the atmosphere. The eruption 
of Tambora volcano on Sumbawa Island in Indonesia in April 1815 was the largest ash eruption 
in recent historic times, producing a bulk volume of about 150 km’ of pumice and ash (Stothers 
1984). The loss of life and the destruction of agricultural land on Sumbawa and neighbouring 
Lombok were catastrophic. In the aftermath of the Tambora eruption, in order to obtain more 
information about the effects on Java and the surrounding islands, the Lieutenant Governor of 
Java, Thomas Stamford Raffles, circulated a letter with three brief questions. The following 
questionnaire was completed by the Resident of Surakarta in eastern Java describing local 
eyewitness accounts of the effects of the Tambora eruption (catalogued in Blagden 1916). It gives 
a vivid picture of the effects of the massive eruption some 800 km from the volcano. (The style, 
punctuation, and spelling of the original handwritten report in the MacKenzie Collection of the 
British Library has been retained throughout). 

Questionnaire and Response by the Resident of Surakarta 

Points of Enquiry 

Circular of the Honble [T.S. Raffles] the Lieut Governor [of Java] 

First, the effects of the eruption of Sumbawa in April 1815 would appear to have been first 
noticed at Banjuwangie on the Ist and at Batavia on the 6th of April but the atmosphere 
would appear to have been successively affected by the ashes between the 10th and 14th. On 
what day and at what hour were they first noticed in different parts of your Residency - 
how long - when did they continue and what was the nature of them? 

At Souracarta the first explosions were heard on Wednesday the Sth of April between the hours 
of 4 and 6 PM, distinct and separate sounds exceeding the number of twenty were perceived with 
irregular intervals greatly resembling a military operation, but more that is denominated mortar 
practice than a regular cannonade. On the successive evenings of the 6th, 7th, 8th and 9th, 
occasional noises were heard which were mistaken for distant thunder. During these days the 
opacity of the atmosphere, resembling former volcanic eruptions on this Island, first indicated 

' Earth Systems Group, Department of Applied Sciences, New York University, New York, New York 10003, 
U.S.A. Also at NASA Goddard Space Flight Center, Institute for Space Studies, New York, New York 10025, 

the probable cause of the explosions which, by a person unaccustomed to their effects could not 
be distinguished from the reports of guns or thunder etc. 

On Monday the 10th, a very slight fall of dust was perceived, but alone by the most attentive 
observation, and the explosions continued at intervals in the east. 

On Tuesday the 11th the reports were more frequent and violent through the whole day: one of 
the most powerful occurred in the afternoon about 2 O’Clock, this was succeeded, for nearly an 
hour by a tremulous motion of the earth, distinctly indicated by the tremor of large window 
frames; another comparatively violent explosion occurred late in the afternoon, but the fall of dust 
was scarcely perceptible. The atmosphere appeared to be loaded with a thick vapour: the Sun was 
rarely visible, and only at short intervals appearing very obscurely behind a semitransparent 

The day on which the opacity of the atmosphere first commenced had not been noted accurately - 
but its continuance was above twelve days, and even at the commencement of the present month 
it was not entirely dissipated. 

From the Sth to the 18th of the last month the Sun was not distinctly perceived, and if his rays 
occasionally penetrated they appeared as observed through a thick mist. The general darkening 
of the atmosphere was strikingly exhibited by such objects of which the prospect is familiar; thus 
for instance at Souracarta the Mountain above continued invisible through all this period, and 
even near objects were clouded or its appearance obscured by smoke - 

On Wednesday the 12th the appearance of day light showed a very copious discharge of dust, this 
gradually increased till 1 PM and then appeared to diminish but was still very discernible at 
sunset: the following day (the 13th) it was still rarely perceptible and gradually and successively 

On the 12th a considerable darkness was occasioned by the abundance of the fall of dust: every 
operation which required strong light was almost impossible within doors. The gloomy 
appearance caused by the rain of dust "Udshan abu” need not be described as it was uniform in 
every part of this Island to which the discharge extended. It may be remarkable that an unusual 
sensation of chillings was felt during the whole of the 12th this was in great measure (tho’ 
probably not exclusively) occasioned by the temperature: the thermometer at 10 O’Clock AM 
stood at 75 and 1/2 degrees of Fahrenheit Scale. It would appear that the subterranean 
commotion, like the discharge of dust, was propogate or travelled from east to west as the 
explosions were later perceived in the Western parts of the island: it is likewise highly probable 
(which must however be determined by a comparison of various and accurate remarks made in 
different parts of the island) that the most violent explosions were not simultaneous, but that the 
combustion caused locally more powerful shocks in particular parts from Banju-wangie perhaps 
to the western extremity. Something like this was remarked during the combustion of the Kloet 
in 1811, when the explosions were much more violent at Batavia than at Souracarta although the 
latter was much nearer to the burning Mountain. It would appear from creditable information that 
effects were more sensibly felt along the Southern Shore of the Island and that the tremulous 
motion of the earth was there more violent - a very uncommon rising of the water was also 
perceived about the period of the most violent explosions at Harang bollong, Kadilangu etc. but 
the day and hour had not been noted with sufficient accuracy for any decided inference. The 
colour of the dust of the present eruption is ash grey inclining to brown it is a most impalpably 


fine, divided earthy substance, if water is added it diffuses the peculiar odour of clay; it does not 
acquire ductility enough to be moulded, but has been observed to improve the quality of the 
common clay of the Island in the manufacture of pottery. Its chief component parts are Silecious 
and Aluminous earth. It is evidently a finely divided Lava, the iron of which having by means 
of gravity subsided in the vicinity of the Volcano. Scarcely any of the particles are attracted by 
the magnet in this it differs from volcanic dust which was thrown from the Gunung Gunter in 
1803 and, being precipitated about Batavia, possessed a blackish colour and was strongly attracted 
by the magnet. The dust which was exploded by the Gunung Klut in June 1811 differs from the 
present as far as can be determined without chemical analysis only by having a blueish grey 
colour, and in being less finely divided; it was supposed to possess superior qualities for the 
manufacturing of pottery but had not ductility enough to be moulded alone. 

Have any injurious consequences resulted from within your Residency as affecting the 
salubrity of the Country, or in the destruction of the Crops or Cattle respecting the latter, 
state the particulars, if any and in what manner the injury may have been effected. 

If the generality of the discharge of the volcanic dust is considered and the abundance of the 
substance which covered the earth and of vegetation for many days, its effects on the health of 
the animals were inconsiderable: instances of mortality among cattle particularly Buffaloes and 
Cows in this neighbourhood during the continuance of and since the rain of dust are Solitary, and 
leave it doubtful whether they must be ascribed to this or other accidental cause. In a few cases 
(within my observation) death was induced suddenly: these may probably be ascribed to this 
cause, but the inquiries I have made have confirmed the opinion that the health of the Cattle has 
not been (in a general manner) injuriously affected. It should be kept in view in determining the 
question, that previously to the rain of dust the Buffaloes in particular districts were affected by 
an epidemic disease denominated Puttie by the Natives of which several died and the mortality 
has in some degree continued to the present time. Neither Horses, Sheep, or Goats have been 
sensibly affected. 

An injury of a more serious nature threatened the crops of rice - but the forward state of 
cultivation has preserved this grain in most of the neighbouring districts and such a season of 
abundance as the present has not been known for many years: it has been observed by various 
persons who are conversant with the cultivation of this grain, that plantations in which the rice 
had nearly acquired maturity were not affected, but the dust falling upon the grain newly 
transplanted in many cases destroyed the young plants. This is in some degree rendered probable 
by the nature of the volcanic substance, and its effects would be more powerful towards the 
period of the terminations of the rains or where a deficiency of moisture prevailed. Falling upon 
the young plants and fields sparingly supplied with water it would from its clayey nature absorb 
their juices and destroy them. 

What was the general opinion at the time regarding the locality of the volcano? 

The general opinion at this place ascribed the eruptions to the Mountain Klut of which three 
previously similar "rains of ashes" were recollected by all aged inhabitants. 


The above report documents that Tambora was quite active for at least six days prior to the 
cataclysmic eruption of 11 April 1815. Note that the eruption was misidentified with Klut (Kelut) 


volcano in Java during the ash rain. The reply gives evidence of a direct cooling associated with 
the ash cloud, and such a cooling effect was observed as far away as Madras, India, where 
midday temperatures fell below freezing as the cloud passed overhead (Stothers 1984). The 
anomalous weather of the infamous summer of 1816 was quite likely related to the radiative 
perturbation by stratospheric H,SO, aerosols generated by the eruption. Without doubt, a similar 
eruption in Indonesia today would be a regional disaster, and would create a global atmospheric 
perturbation of a magnitude not seen in almost two centuries. 


A grant from the American Philosophical Society supported a literature search at the British 
Library for information pertaining to the aftermath of large volcanic eruptions in the nineteenth 
century. The author thanks I.A. Baxter of the India Office Library, Blackfriars Road, London, 
for his help, and S. Self, H. Sigurdsson, and R.B. Stothers for valuable discussions. This is a 
slightly altered version of a paper published by the author in EOS 1989, p. 1559, (copyright by 
the American Geophysical Union). 


Mackenzie Collection: Private, Document 2:33, pp. 193-198, 1916. In: C.A. Blagden, 
Catalogue of Manuscripts in European Languages belonging to the Library of the India 
Office, Volume 1: The Mackenzie Collection, Part I: the 1822 Collection and Private 
Collection, p. 43. Oxford University Press, London. 

Stothers, R.B. 1984. The great Tambora eruption in 1815 and its aftermath. Science 224:1191- 

The Eruption of Tambora in 1815: Environmental Effects and Eruption - 

Haraldur Sigurdsson’ and Steven Carey’ 


New studies of deposits from the 1815 eruption of Tambora volcano provide data on eruption 
dynamics, mass eruption rate and volcanic volatile emission to the atmosphere. These data form 
a basis for assessment of the environmental impact of the eruption. Initial phases of activity were — 
two plinian explosive eruptions on 5 and 10 April with column heights of 33 and 43 km, and 
mass eruption rate of 1.1x10° and 2.8x10* kg/s respectively. The calculated column heights 
therefore indicate a major injection of volcanic ash and volatile gases to the stratosphere during 
the eruption. Rapid transition to pyroclastic flow generation occurred late on 10 April, when the 
bulk of the material was erupted at a rate of 5.4x10* kg/s, producing widespread co-ignimbrite 
ash fall. A large component of the co-ignimbrite ash fall was produced by explosive interaction 
of hot pyroclastic flows and sea water, when flows advanced into the ocean around Tambora. 
Total erupted mass is estimated as 50 km’ dense-rock equivalent, or 1.4x10'* kg. Petrologic 
estimates of volatile yield to the atmosphere during the eruption indicate that sulphur degassing 
formed a stratospheric aerosol mass equivalent to 1.75x10"' kg sulphuric acid, in agreement with 
volcanic aerosol estimates based on ice-core evidence. Furthermore, volcanic degassing of 10" 
kg HC@é and 7.4x10" kg HF occurred, but the fate of these species in the atmosphere is 
unknown. Climatological data indicate a short-term northern hemisphere surface temperature 
decrease of 0.7°C following the eruption, and this climatic response agrees with the empirical 
relationship observed between sulphuric acid volcanic aerosol mass and temperature decline 
observed after several major explosive volcanic events. It is likely, however, that the observed 
surface temperature decline is not solely due to the Tambora event, as a cooling trend was 
already in progress prior to the eruption. 


The 1815 Tambora eruption on the island of Sumbawa in Indonesia exceeded in magnitude any 
other volcanic eruption in historical times, producing over 50 km’ of magma. As a measure of 
the uniqueness of this great natural disaster, it is remarkable that we have to search some 20,000 
years back in the geological record to find an explosive eruption of greater magnitude: the 
Shikotsu eruption in Japan (Katsui 1959). The volume of material erupted from Tambora is an 
order of magnitude greater than that discharged in the celebrated Krakatau eruption of 1883, and 
two orders of magnitude greater than in the 1980 Mount St. Helens eruption. Locally, 92,000 
people died on Sumbawa and adjacent islands, either directly from effects of the eruption or from 
the ensuing famine and epidemic. In addition to its significance as a geological process, the 
eruption had unprecedented impact on the Earth’s stratosphere. The eruption injected enormous 
quantities of sulphur, chlorine and fluorine gases into the stratosphere, leading to a variety of 
global atmospheric phenomena, and was probably responsible for the marked climatic 
deterioration of 1816. Thus, although the eruption is of great importance to the study of 
volcanology, its greatest scientific significance probably relates to the environmental effects, i.e., 

' Graduate School of Oceanography, University of Rhode Island, Kingston, Rhode Island 02881, U.S.A. 


its impact on the chemistry of the atmosphere and on climate (See also Vupputuri, this volume). 
With the growing realization of connections between the biosphere, atmosphere and geosphere 
and the recognition of global environmental and climatic change brought about by human activity, 
the detailed study of the effects of Nature’s own large-scale experiments such as the Tambora 
eruption can greatly aid in our understanding of short- and long-term changes in the global 

In this paper we summarize our findings based on a new study of the Tambora deposits, 
involving two expeditions to the volcano in 1986 and 1988, which provide fresh data on the 
eruption dynamics and erupted mass. In addition, our recent petrologic study of the sulphur, 
chlorine and fluorine yield of the eruption to the atmosphere gives quantitative estimates of 
degassing and provides a framework for modelling the environmental impact of this great 
volcanic pollution event. 

Chronology of the 1815 Eruption 

Before 1815, Tambora volcano was conical in form, possibly with two peaks, and the highest 
mountain in the Sunda Islands. When sailing east from Java, Tambora appeared equally 
prominent on the horizon as the 3,726-m high Rinjani volcano on Lombok, and Zollinger (1855) 
estimates that the volcano may have been over 4,000 m before the eruption. His estimates are 
based on discussions with people in Sumbawa, who maintained that the volcano had lost at least 
one third of its height. The maximum height of the caldera rim after the eruption is 2,850 m. 

Contemporary local sources about the 1815 Tambora eruption are mainly newspapers and 
government accounts - particularly the Asiatic Journal. These accounts are especially useful for 
establishing the timing of various eruptive phases, and the extent and nature of their effects. This 
section summarizes important eyewitness observations that are relevant to interpreting the 
pyroclastic deposits studied in the field (Rampino, this volume). 

More than three years before the great eruption, a thick cloud had formed over the peak, which 
not even the strongest winds could dissipate (Zollinger 1855). It gradually grew darker and 
larger, and extended farther down the volcano’s flanks. Explosions were heard from the volcano 
during this time; first only a few and weak, but gradually they became more frequent and louder. 
People living around the volcano sent delegates to the government authorities in Bima on 
Sumbawa requesting an investigation of these phenomena. The authorities sent a man by the name 
of Israel, whose brother was still alive at the time of Zollinger’s visit. Israel reached the Tambora 
region on the evening of 9 April, the day before the climax of the eruption, and was killed during 
the activity the following day. 

On the evening of 5 April, the first major eruption began and was heard widely in the Indonesian 
region. The explosions heard in Java resembled cannonfire and soldiers in Yogyakarta (central 
Java) combed the land and seas for invaders (Figure 1). Ash fell "like fine snow" in Banjuwangi 
in eastern Java, accumulating up to one-half inch (1.3 cm) thickness. Minor ash fall also occurred 
at Besuki in east Java. At Solo (central-eastern Java, 800 km from Tambora) sounds of explosions 
commenced on the evening of 5 April. The naval vessel Benares was in Macassar on 5 April, 
about 350 km NNE of Tambora (Figure 1). Loud explosions were heard from the south, which 
continued the entire afternoon. At sunset the explosions grew louder and seemed closer. Listeners 
suspected that a naval battle was taking place nearby and sent troops to search the region. 


11 April AM 


300 km Yogyakarta 


10 April PM 
Scale P 


Figure 1: The Indonesian region, showing areas where sounds were heard from the 1815 Tambora 

Ash fell in east Java during the morning of 6 April, but only a trace descended on western Java. 
The sky gradually cleared during the day, but the air was hot and the atmosphere unusually still. 
Between 6 and 8 a.m. on 6 April, loud noises were heard at Ternate, where the ship Teignmouth 
lay at anchor some 1,400 km northeast of Tambora. 

After four days of minor activity, the volcano became very active again on 10 April. Witnessing 
the activity from Sanggar on the eastern slopes of Tambora, the Rajah described the second and 
larger major eruption. At about 7 p.m. three columns of fire rose high from Tambora’s crater, 
uniting in a single firestorm over the volcano. Moments later the entire mountain was a sea of 
glowing flows, which spread in all directions. Large quantities of ash and stones fell on Sanggar 
(Figure 2), “up to two fists in size", but most were no larger than a nut. Between 9 and 10 p.m. 
the ash fall increased, and a strong "whirlwind" descended carrying off houses in Sanggar and 
nearby villages. In the part of Sanggar nearest to Tambora, the largest trees were uprooted by 
the windstorm, and carried off with houses, people and livestock. These descriptions are 
consistent with the passage of a pyroclastic surge through the village. Sea level rose suddenly 
12 feet Gv m): 

At Bima, 80 km east of the volcano (Figure 2), the explosions sounded like heavy mortar fire 
during the night of 10 to 11 April. The town was in complete darkness from the ash cloud 
overhead from 7 a.m. on 11 April to 14 April. Ash fall was so heavy, that roofs of most houses 
collapsed. The air was completely still and there was no wind at sea, but nevertheless the waves 
were very high and flooded the coast and into the town. All boats were torn from their moorings 
and tossed ashore. 


Flores Sea 

Satonda — Kawinda 

Sumbawa Doro Petie 

Gulf of Saleh 

Figure 2: Sumbawa Island, showing Tambora volcano and other places referred to in the text. 

Beginning on 10 April, thunderous noises were heard in many parts of Java, which were much 
louder than on 5 April, especially east of Cirebon (western Java, 1,050 km from Tambora). At 
Banjuwangi in east Java (400 km distant) the evening noises were very loud and shook the earth. 
The sounds became somewhat weaker toward morning the next day, but continued until 14 April. 
At Sumenep (Madura Island, 470 km distant) the noises were like rapid cannonfire. The sky was 
completely obscured by ash and, in some districts such as Solo and Rembang (central Java, 
Figure 1), earth vibrations were felt. 

During the night of 10 to 11 April the Benares reported from Macassar that explosions began 

again and grew in frequency the next morning, shaking both houses and ships. Lightning flashes 

were common and the sky was very dark, especially to the south and southwest. The sea rose 
rom five to seven feet (1.6 to 2.2 m) above normal in Besuki, eastern Java, on the night of 10 

On 11 April the continuing activity was so severe that houses shook in eastern districts of Java. 
The coast of Bali was totally invisible from Banjuwangi in eastern Java, where candles were lit 
at 1 p.m. By 4 p.m. it was pitch-dark and remained dark until 2 p.m. the next day. In Sumenep 
(Madura) the light was so faint that candles had to be lit before 4 p.m. The following night was 
indescribably dark. At about 7 p.m. a tidal wave struck Sumenep Bay raising sea level about four 
feet (1.2 m) for several minutes. Major ash fall also began in Besuki, eastern Java, on 11 April, 
with darkness extending from 4 p.m. on 11 April until 2 p.m. on 12 April. Explosions were also 
heard on 11 April at Ambon (Asiatic Journal, February 1816, p. 116). 


A boat sailing from Timor in the east noted that the sky became very dark as they approached 
Tambora on 11 April. When they were off Tambora, the base of the volcano was engulfed in 
flames and the peak was shrouded in a dark cloud, with fires and flames shooting out. They went 
ashore for water in Sumbawa and found that all boats had been cast ashore by tidal waves. They 
came across a large number of corpses. As they sailed from Sumbawa, they encountered large 
rafts of pumice, which formed thick layers on the ocean hindering their passage. Some pumice 
rafts were so thick that they resembled sandbanks or low cliffs. They were caught in a pumice 
raft over two feet (0.6 m) thick the entire night of 12 April. The vessel Dispatch heard explosions 
on the night of 11 April, when about 7° east of Bima (about 750 km). Rafts of pumice and 
timber were so thick along the coast of Flores, that the ship had great difficulty in making way. 

Effects of the eruption were noted as far west as Sumatra. On the morning of 11 April, loud 
noises were heard at Bengkulu on the south coast of Sumatra about 1,800 km west of Tambora, 
and as far as Terumon in western Sumatra some 2,600 km WNW of Tambora. Explosions were 
also heard on Bangka Island (1,500 km) off the northeastern coast of Sumatra. People from the 
interior of Sumatra reported that the leaves of trees and crops were covered with a layer of very 
fine ash (Asiatic Journal, June 1816, p. 600 and August 1816, p. 164). 

On 12 April only very faint daylight was visible in eastern Java, and objects were barely visible 
at a distance of 100 paces in Solo. Some light returned in Banjuwangi about 2 p.m., but the sun 
was not visible until 14 April. It was unusually cold during this period. Ashfall in Banjuwangi 
was nine inches (of which eight inches (20.3 cm) had accumulated by 12 April), two inches 
(5 cm) in Sumenep and somewhat less in Gresik. West of Samarang (central Java) the daylight 
was little affected. 

At 8 a.m. on 12 April it was dark on the Benares in Macassar, and by 10 a.m. it was so dark 
that nearby ships could not be seen. By 11 a.m. the sky was completely dark, except for a small 
clearing in the east. The ash fell as heavily as snow, and the sea and air were still. By 12 noon 
the faint light in the east had vanished and it was so dark that a hand held in front of the face 
could not be seen. Ash fell all night and was so fine that it penetrated all parts of the ship below 
decks. By 13 April the intensity of the eruption had decreased but its effects were still 
widespread. At 6 a.m. it was still totally dark in Macassar but faint light returned at 7:30 a.m., 
and by 8 a.m. one could discern objects. Sounds of the explosion ceased the following day at 
Banjuwangi but ash fall continued in Macassar, accompanied by a calm and great heat until 
15 April. 

Not until 17 April did the ash fall cease, and heavy rains spread over the region. In Banjuwangi, 
many houses had collapsed under the weight of the ash, and fever and epidemics had broken out 
in several regions affected by the ash fall. On Java the damage to livestock and agriculture was 
most severe in the eastern district around Banjuwangi, where the destruction of crops and grazing 
areas was SO extensive that many horses and cattle died of hunger. 

The devastating effects of the eruption on the local population were first realized when ships 
reached ports on Sumbawa. Benares reached the coast of Sumbawa on 18 April and was trapped 
in large pumice rafts on the sea. The rafts were so large, that they at first took them for 
sandbanks or new islands: they were often over a nautical mile in length, and had varied surface 
features. Large numbers of carbonized and splintered trees were trapped in the rafts. Benares 
dropped anchor at Bima on 19 April, where ash fall was 3% inches (9.5 cm) thick. The harbour 


had changed, and they found eight fathoms (14.9 m) where the depth had been six fathoms 
(11.2 m) before the eruption. 

Some pumice rafts were up to three miles (4.8 km) long, and were still troublesome to navigation 
between Moyo and Sanggar three years after the eruption. Pumice rafts from the volcano drifted 
widely over the southern seas in the following months. Between 1 and 3 October 1815, the ship 
Fairlie, in the Indian Ocean on passage to Calcutta, sailed for two days through extensive pumice 
rafts, about 3,600 km west of Tambora (Asiatic Journal, August 1816, p. 161). These rafts 
travelled at a rate of 0.2 m/s from Tambora and were most likely transported in the South 
Equatorial Current, driven by the southeast trade winds. Ash fall from the eruption also reached 
Brunei in Borneo, where the phenomenon so impressed the local people, that they subsequently 
counted the years from "the great fall of ashes" (Reclus 1871). 

On 22 April, the Dispatch arrived in Bima. It had first dropped anchor near Sanggar, where the 
Rajah had told them that all the land was now a desert and all crops and fruits were destroyed. 
Sanggar Bay was covered with pumice rafts including large trees and remains of houses carried 
out to sea by the eruption. The volcano was still covered in dense clouds of ash and steam. 
Smoke emanated in many places from hot flows of ash on the lower flanks, which had also 
entered the sea. 

The British Governor of Java sent Lieutenant Owen Philipps to Sumbawa to study the event and 
its effects on the people. On the way from Bima to Dompu (Figure 2), Philipps observed a large 
number of corpses along the road. Villages were abandoned and houses had generally collapsed 
under the weight of the ashfall. The few survivors wandered about in search of food. The 
population had been affected by severe diarrhoea, which had caused many deaths. The people 
blamed this on their drinking water, which was contaminated with the volcanic ash. Horses and 
other livestock were also killed in large numbers by this disease. The Rajah of Sanggar met with 
Philipps in Dompu. The misery of his people was much worse than in Dompu and even one of 
the Rajah’s daughters had died of hunger. Coconuts were the only food supply of the ruined 
village, where starvation was severe. Philipps gave him some rice, for which the Rajah gave 
thanks with tears in his eyes. 

Zollinger (1855) describes the misery of the remaining population. Many continued to wander 
in search of food and willingly sold themselves as slaves, sometimes for a few pounds of rice. 
His studies indicate that about 10,100 people died in Sumbawa directly by the effects of the 
eruption, most likely in pyroclastic flows and surges (Table 1). Contemporary estimates of 
number of fatalities in several villages vary. Thus, for example, Tobias claims there were 10,000 
deaths in Tambora village alone, whereas Philipps claims 12,000 victims in this village. In 
addition, 37,825 died by starvation and 36,275 migrated from Sumbawa. Zollinger estimates that 
at least 10,000 died in Lombok from starvation and disease, but the loss there was much more 
severe according to Van der Broeck (1834), who states that the population of Lombok was 
reduced from 200,000 to 20,000 by the effects of the eruption. Zollinger claims his numbers are 
all minimum estimates. Junghuhn (1850) estimates that the fatalities on Sumbawa were 12,000 
and that 44,000 died on Lombok, but his estimate does not include the starvation victims on 
Sumbawa. The most-quoted fatality figures of the eruption are those of Petroeschevsky (1949), 
who estimates that the total number of victims was 92,000 - 48,000 on Sumbawa and 44,000 on 
Lombok, or 35 and 22.5% of the estimated total population of these islands, respectively. 


Table 1: Fate of the Human Population in Sumbawa.' 

' After Zollinger (1855). 

The only village near Tambora that remained undamaged was Tempo, with 40 inhabitants. Of 
the total population of 12,000 of Tambora and Pekat, only five or six survived. All trees and 
vegetation north and west of the volcano were completely destroyed, with the exception of a high 
point near the village of Tambora. Zollinger remarked on the long-term effects of the eruption 
on Sumbawa’s climate and vegetation. Soil became very dry, rainfall decreased, and all 
vegetation suffered a severe setback, and would take an estimated several hundred years to 
recover fully. 

Pyroclastic Deposits from the 1815 Eruption 

As a consequence of the eruption, the upper part of the volcano collapsed to form a 6-km 
diameter, 1,200-m deep caldera with a total volume of about 28 km?, and Tambora lost about 
1,200 to 1,400 m of its height, corresponding to about 6 km’ or a total of 34 km*. The void 
formed by the caldera collapse represents in part rock formations ejected from the volcano, and 
in part the subsidence of the volcano’s edifice into the underlying magma chamber. The former 
can be evaluated from proportion of lithics in the fall deposits, which is about 5.5 wt.% 
(Sigurdsson and Carey in press, Table 2) or less than 4 km’ of rock. Ejection of solid rock can 
consequently account for one-tenth of the caldera volume. The total ejected mass of magma is 
1.3x10"* kg, less the lithics, corresponding to about 50 km’ of magma withdrawn from the 
reservoir - substantially larger than the observed caldera volume. Subsidence into the emptying 
magma reservoir is regarded as the dominant mechanism of caldera formation. 

The deposits laid down outside the caldera during the eruption reflect two major processes: 
(1) early explosive activity (plinian and phreatomagmatic) producing high eruption columns and 
four tephra or ash-fall deposits; and (2) subsequent ignimbrite phase activity during collapse of 
the eruption column, producing at least seven pyroclastic flows and surge deposits, with 
associated large-volume co-ignimbrite ash falls. 


Table 2: Tambora 1815; Composition of Glass Inclusions in Plagioclase Phenocrysts.' 



Number of 

Number of 

Volatiles by 

Water by 

Sulphur (ppm) 
Chlorine (ppm) 
Fluorine (ppm) 

57.37 (1.28) 
0.6 (.07) 
19.66 ( .41) 
4.56 ( .39) 
0.27 ( .05) 
6.19 ( .25) 
5.09 ( .47) 
0.06 ( 0) 





57.01 (.51) 
0.56 (.06) 
19.58 (.22) 
4.47 (.20) 
0.28 (.08) 
1.09 (.07) 
3.09 (.08) 
5.5 (.50) 
5.59 (.27) 
0.04 (.02) 



589 +94 

2,057 732 

56.37 (.29) 
0.73 (.01) 
19.43 (.06) 
4.66 (.07) 
0.25 (.07) 
1.18 (.03) 
2.87 (.01) 
S89, 1. 
5.69 (.09) 
0.31 (.13) 





56.88 (1.24) 
0.60 ( .10) 
19.88 ( .45) 
4.73 ( .48) 
0.19 ( .08) 
1.37 ( .15) 
2.85 ( .25) 
644 ens 
5.35 ( .86) 
0.36 ( .15) 





2,817 +1253 

56.58 (1.13) 


20.17 ( .21) 

5.23 ( .94) 
0.24 ( .05) 
1.75 ( .44) 
2.50 ( .25) 
3.15 ( .54)? 
6.00 ( .48) 


381+ 44 
2,106 + 163 
1,185+ 87 

1 - glass inclusions in plagioclase from plinian fall layer F-2, sample TB-42; 2 - glass inclusions in plagioclase from 
lower part (0 to 5 cm) of plinian fall F-4, sample TB-86; 3 - glass inclusions in plagioclase from upper part (10 
to 15 cm) of plinian fall F-4, sample TB-88; 4 - glass inclusions in plagioclase from co-ignimbrite fall deposit F-5, 

sample TB-136; 5 - glass inclusions in plagioclase from tephra fall, sample T58-A (Devine ef al. 1984). 

(0.37 wt.%). 

The four initial explosive events produced widespread tephra fall deposits, which can be traced 
at least to Lombok, 150 km west of the caldera (Sigurdsson and Carey in press). The basal F-1 
ash fall is the product of phreatomagmatic explosions, resulting from interaction of magma with 
the hydrothermal system of the volcano (Figure 3). Historical evidence (Petroeschevsky 1949) 
indicates that the volcano was mildly active in the period 1812-15, when "rumblings and dense 


Not corrected for sodium loss during microprobe analysis. All other values are corrected. 

Water by difference is calculated as volatiles by difference minus S, Cf and F 

clouds" were noted. The F-1 ash fall probably originated during this early activity, as magma was 
making its way from a deep reservoir toward the surface and periodically erupting in small 
outbursts. The total volume of tephra erupted during the phase of activity was about 0.1 km’. 
Evidence from our excavations in the ruins of the ancient Tambora village, 2 km east of Tambora 
Coffee Estate, indicates that the F-1 event took place long before the subsequent activity. The F-1 
layer is absent from the village, indicating its complete erosion before the 5 April eruption. 

Stratigraphy of 1815 Tambora Deposits 

PF-| Pyroclastic flow April 10-1 
(1-4 meters) 

Pyroclastic surge April 10 

Vea S772 
Sey i 




Lh é - 

“wv, ¥, 
C3 °, 



F-4 Plinian Pumice Fall April 10 


F-3 Phreatomagatic Ash Fall April 5-10 




F-2 Plinian Pumice Fall April 5, 1815 

Phreatomagmatic Ash Fall Pre-April 5, 1815 

Figure 3: Stratigraphy of the 1815 pyroclastic deposits in a typical section at Gambah on the 
northwestern slopes of Tambora volcano, 25 km from the caldera. Dates on the right of the 
stratigraphic column indicate the timing of successive eruptive phases, based on historical 

The F-2 pumice fall layer marks the first major explosive eruption during 1815. The distribution, 
lithology and grain-size of this deposit indicate typical plinian activity. About 1.2 km* of material 
was ejected. We correlate this plinian eruption with the explosion of 5 April 1815 that was heard 
in Jakarta (1,250 km away) and Ternate (1,400 km away) and caused ash fall as far as Besoeki 
in East Java (Raffles 1835). 


After the F-2 plinian event, Tambora lapsed into a state of low-level activity from 5 to 10 April. 
During this period, several smaller explosions produced tephra fall, which forms layer F-3 
(Figure 3). The deposit is highly fragmented, like the first layer of the eruption (F-1), but the 
evidence of the phreatomagmatic activity is not as compelling. 

A second major plinian eruption produced the F-4 pumice fall layer. Grading of this deposit 
shows a rapid rise of the eruption column during the first third of the eruption, followed by a 
slow decline. The F-4 layer is much thicker and coarser than the earlier F-2 plinian fall, although 
similar in lithology. Despite its high intensity, this phase of the eruption ejected only a moderate 
amount of material (3 km* of tephra). The F-4 fall deposit is clearly from the beginning of the 
10 April paroxysmal event. 

The Rajah of Sanggar reported an intensification of activity at about 7 p.m. on 10 April, followed 
by a rain of pumice on Sanggar, east of the volcano, at approximately 8 p.m. Tephra fall 
continued until about 10 p.m. when the village experienced winds that uprooted trees and 
buildings. The whole volcano appeared like a flowing mass of "liquid fire". This event is marked 
clearly in the volcanic stratigraphy everywhere on Sanggar Peninsula by the abrupt transition 
from F-4 plinian pumice fall to the overlying charcoal-bearing surge and pyroclastic flows (Figure 
3). The change in the eruption mechanism may have been primarily due to continued vent erosion 
during F-4 plinian activity, leading to eruption column collapse, with resulting pyroclastic flows 
and surges. No significant time break may have occurred during the transition. 

In distal localities the F-4 plinian fall is overlain by a 12- to 25-cm thick greyish-brown, poorly- 
sorted, silty-sandy ash (F-5). Unlike the other fall deposits from the 1815 eruption that show 
systematic thinning with distance from source, the F-S ash fall retains a constant thickness to a 
remarkable degree - in fact thickening appreciably to the west, away from the volcano. 
Consequently the F-5 layer represents an increasing proportion of the total fall with distance from 
source, increasing from about 25% of the total fall deposit thickness at 40 km, to about 80% 
beyond 90 km (Figure 4). 

The F-5 layer does not correspond to any fall deposits in the proximal area, but is 
Stratigraphically equivalent to the surges and pyroclastic flows. We therefore consider the F-S 
deposit formed primarily from ash and pumice fallout from an eruption column generated during 
the surge and pyroclastic flow phase, i.e., a co-ignimbrite and co-surge ash fall deposit. The co- 
ignimbrite ash fall was not only generated by glass elutriation from the convecting eruption 
columns and flows, but also by wholesale depletion of the fine fraction of crystals and glass alike 
from the column and flows. We attribute this depletion to explosive interaction between 
pyroclastic flows and the sea along the coast of the Sanggar Peninsula (Sigurdsson and Carey in 
press), based on comparative grain-size studies of inland and coastal pyroclastic-flow deposits. 
Our model proposes the creation of large secondary eruption columns around the peninsula of the 
volcano, where high-temperature pyroclastic flows were discharged into, and reacted explosively 
with seawater. The secondary plumes consisted mostly of fine-grained (< 200 micron) ash and 

Total Erupted Mass 
It is generally recognized that the Tambora eruption involved an exceptionally large volume of 
magma, although quantitative estimates have varied greatly. Thus, Zollinger (1855) estimated the 

ash fall volume at >1,000 km®, Junghun (1850) 318 km°, Verbeek (1885) 150 km*, Sapper 
(1917) 140 km?, Pannekoek van Rheden (1918) 30 km’, and Petroeschevsky (1949) estimated 


total ash fall of 100 km? on the basis of observed thicknesses. A reassessment of the ash fall 
volume by Stothers (1984) led to an estimate of 150 km’, and Self er al. (1984) estimate 175 km’. 
New estimates can now be made on basis of our recent field work, Sigurdsson and Carey (in 


‘3 : 
Ew E 
a 3 
E = 
‘= 10 ° 



0 t 
0 40 80 120 160 0 40 80 120 160 
Distance from Source (km) Distance from Source (km) 

Xtal/Glass Ratio 

~0 40 80 120) "160 
Distance from Source (km) 

Figure 4; Characteristics of the F-5 co-ignimbrite tephra fall deposit as a function of distance from 
source, showing: (a) variation in thickness in cm; (b) thickness of F-5 co-ignimbrite ash fall 
as a fraction of total ash fall thickness; (c) crystal/glass ratio of the co-ignimbrite ash fall. 
Horizontal line in (c) is the crystal/glass ratio in the erupted magma, as determined in 
artificially-crushed pumices from the pyroclastic flows. 

As shown above, the products of the eruption form a multi-layer deposit, reflecting several 
processes in action. The four early fall deposits produced during activity from 5 to 10 April (F-1 
to F-4), have a total volume of 4.6 km’, corresponding to 1.8 km’® of dense rock, or about 
4.3x10" kg of magma. While this represents an eruption larger than the 1980 Mount St. Helens 
event, and comparable in volume to the 1982 El Chichén event, these early April falls from 
Tambora represent only 5% of the total erupted mass in 1815. The co-ignimbrite ash-fall layer 


F-5 is the dominant part of the deposit. Volume of the total ash fall can be estimated from 
contemporary accounts of ash fall and thickness, and deep-sea core evidence (Neeb 1943). Using 
the isopach map compiled by Self et al. (1984) and shown in Figure 7, we estimate that 90 km’ 
of tephra was deposited within the 1 micron isopach, corresponding to 22 km’ dense-rock 
equivalent of distal fall (5.5x10"° kg). The F-5 fall must represent about 92% of this volume, as 
the combined volume of the earlier F-1 to F-4 fall layers is only 1.8 km® DRE. An estimate of 
density of the deposit is required in order to assess the erupted mass. During the eruption, ash 
fell on decks of the ship Benares near Macassar in Sulawesi. A pint of the ash was reported to 
weigh 12% oz, corresponding to a deposit density of 611 kg/m® of the fresh-fallen ash (Asiatic 
Journal 2, 1816, p. 166). A minimum mass of 5.8x10" kg is therefore represented by the fall 

Studies of the volcano and its deposits indicate that a large mass of pyroclastic flows entered the 
ocean during the eruption (Sigurdsson and Carey in press). We estimate a pyroclastic flow deposit 
volume of about 30 km’, equivalent to 8.2x10" kg. Thus, the total erupted mass is of the order 
1.4x10'* kg of magma. No historical eruptions have produced as large a mass of magma as 
Tambora, which emitted more than twice the mass of the nearest large-magnitude event, i.e., the 
1783 Laki eruption in Iceland. 

Mass Eruption Rate and Column Height 

Pumice and lithic isopleth maps of the F-2 and F-4 layers are presented by Sigurdsson and Carey 
(in press). The area encompassed by a specific isopleth is considerably larger for the F-4 layer, 
demonstrating the greater intensity and dispersal of that plinian event. The distribution of the two 
Tambora plinian layers is compared with several other well-documented plinian fall deposits 
(Figure 5). Our new isopleth data indicate that the two Tambora plinian fall deposits had greater 
dispersal than any plinian eruption in historic times. Despite the fact that their dispersal compares 
with some of the largest known plinian fall deposits in the geological record, the thicknesses and 
thus volumes of the two Tambora fall deposits are relatively small (Figure 5). The great dispersal 
of clasts during the two plinian eruptions of Tambora is noteworthy and has important 
implications for existing models of the 1815 eruption. 

The dispersal characteristics of the pumice fall preserve information about the dynamics of the 
eruption column and the atmospherically-dispersed plume. Thus, the geometry of lithic isopleths 
can be used to determine the maximum eruption-column height and average wind speed for a 
specific fall layer (Carey and Sparks 1986). The half-width of an isopleth measured perpendicular 
to the main dispersal axis is primarily a function of the eruption column height, whereas the 
maximum downwind range along the axis is controlled by both column height and average wind 
speed. Data from the 3.2-cm diameter lithic isopleths of the F-2 and F-4 layers indicate eruption- 
column heights of 33 and 43 km, respectively (Figure 6). This places the F-2 column higher than 
the maximum height achieved by the 79 A.D. plinian eruption of Vesuvius (Carey and Sigurdsson 
1987), and the F-4 column is slightly higher than the great 1956 Bezyminanny eruption 
(Gorshkov 1959). The F-4 plinian phase is thus the most energetic plinian activity ever recorded 
in historic times, and is exceeded in intensity by only one eruption in the geological record - the 
“ultraplinian" Taupo pumice fall in New Zealand (Walker 1980). 


Isopleth Area (km2) 

Figure 5: 


Figure 6: 

Thickness (cm) 

0 20 40 60 80 
Lithic Diameter (mm) Distance from Source (km) 
Comparison of 1815 Tambora fall deposits with characteristics of deposits from other major 
volcanic eruptions. (a) Plot of lithic isopleth area versus lithic diameter for the F-2 (5 April) 
and F-4 (10 April) plinian fall deposits compared with the plinian falls from the eruptions of 
Vesuvius, Italy (79 AD), Osumi, Japan, and Taupo and Waimihia, New Zealand. (b) Plot of 
thickness versus distance from source for the F-2 and F-4 Tambora plinian fall layers 
compared to other well-known plinian deposits, as in Figure 5 (a). Note that despite the fact 
that Tambora layers are very widely dispersed, they are substantially thinner than other major 
plinian fall deposits. 
20 \? \? \2 \2 
F-4 .@ 
3.2cm DIAMETER OF Ooo mene 
40 km 
, % Tambora I815 
10 30 km 
= 0 O Nevadodel Ruiz 
O El Chichon 1982 
5 20 km 4 Santa Maria 1902 
® @ Fogo 1563 
4 Tarawera |886 
lO km ® Askja 1875 
O 5 10 15 20 25 30 
Plot of isopleth half-width versus maximum downwind range of the 3.2-cm diameter lithics 

isopleths for the F-2 and F-4 plinian Tambora fall deposits, compared with other well-known 
plinian falls. Diagonal lines are wind-velocity contours in m/second, and horizontal lines are 
maximum eruption-column heights in kilometres. Note the 43 km high F-4 plinian column 
from 10 April 1815 above Tambora volcano, and the 33 km high column from the F-2 
eruption on 5 April. 


The estimates of eruption-column height can be used to calculate the eruption rate by using 
relations for a tropical atmosphere (Sparks 1986). Our calculations indicate a rate of 1.1x10* kg/s 
for the F-2 phase and 2.8x10* kg/s for the more energetic F-4 event. With these values it is 
possible to estimate the duration of the events by simply dividing the total mass of tephra in each 
layer by the perspective rate of magma discharge. Assuming that the maximum magma-discharge 
rate was active throughout the plinian eruptions, each layer would have been ejected in 2.8 hours. 

Duration of the co-ignimbrite fall was about three days, judging from the historical reports, e.g., 
from Madura Island, 500 km WNW of the volcano (Figure 7). This is the period of the 
sedimentation of tephra from the atmosphere and thus represents the maximum duration of the 
eruption which began on 10 April. In order to accommodate the total ignimbrite and co- 
ignimbrite mass in this period (1.4x10" kg), we infer a minimum ignimbrite mass-eruption rate 
of the order of 5.4x10* kg/s, or about three times the peak rate during the eruption of the 
preceding F-4 plinian fall. This is about half the rate of the highest intensity event known: the 
Taupo eruption in 130 AD, with an eruption rate of 1.1x10’ kg/s (Walker 1980). 





Lombok Sumba 

+ Snellius Cores 

Figure 7: Isopach map of the total ash fall from the 1815 Tambora eruption, based on contemporary 
reports of ash fall and evidence from bottom samples collected during the Snellius Expedition. 


Volatile Emission from the Tambora Eruption 

Recent studies have shown that quantitative estimates can be made by petrological methods of the 
mass and type of volatiles (e.g., sulphur, chlorine and fluorine) released during volcanic 
eruptions. The potential of trapped glass inclusions as recorders of pre-eruption volatile content 
of magmas was first recognized by Anderson (1974), who applied this method in estimating the 
volcanic volatile contribution to the sulphur and chlorine budget of the oceans. The method was 
also applied in the 1976 St. Augustine eruption by Johnston (1980), who demonstrated the 
potentially great contribution of volcanic eruptions to the chlorine budget of the stratosphere. 
These studies paved the way for the petrologic estimates of volcanic degassing during earlier and 
prehistoric eruptions (Sigurdsson 1982; Devine et al. 1984; Palais and Sigurdsson 1989). When 
compared with other determinations of volatile emission based on ice-core acidity and 
atmospheric observations, the petrologic estimates yield similar results for the same eruptions 
(Sigurdsson et al. 1985). 

In the first petrologic study of volcanic volatiles from the 1815 Tambora eruption, Devine et al. 
(1984) found that seven glass inclusions in feldspar phenocrysts from a single pumice sample 
contained on the average 380 ppm sulphur, 2,100 ppm chlorine, 1,190 ppm fluorine. We have 
analyzed glass inclusions in plagioclase phenocrysts and matrix glasses in five tephra samples 
from the 1815 eruption, representing all major deposits produced during the event. Our results 
(Tables 2, 3) show that the eruption tapped a homogenous body of trachyandesite magma, with 
no systematic chemical gradients. We find that the average pre-eruption concentration of volatiles 
is 570 ppm sulphur, 2,220 ppm chlorine and 1,190 ppm fluorine, whereas the degassed matrix 
glass has on the average 266 ppm sulphur, 1,486 ppm chlorine and 680 ppm fluorine. These 
results indicate, that about 53% of the pre-eruption sulphur content of the magma was lost to the 
atmosphere during the eruption, accompanied by loss of 33% of the chlorine and 43% of the 
fluorine. In addition, the results indicate a pre-eruption water content of about 2 to 2.4 wt.% in 
the magma. 

Table 3: Matrix Glass Composition of Tambora 1815 Tephra (parts per million). 

Sample Numbers 

TB-42 TB-86 TB-88 TB-136 TB-87 T58-A 
Sulphur 363 +57 126+17 241 +67 196+48 362433 309 +7 
Chlorine 1,523+174 1,460+142 1,621+22 1,476+69 1,627+139 1,211+50 
Fluorine - - - - - 679 +69 

' Errors are one standard deviation of the average. 


With a known total erupted mass of magma of 1.4x10'* kg, the minimum mass of volatiles 
emitted to the atmosphere can be estimated from the difference in volatile concentration between 
glass inclusions and matrix glasses. These calculations show that about 4.3x10" kg of sulphur 
were released to the atmosphere, 1x10" kg chlorine, and 7x10" kg fluorine. These improved 
estimates are somewhat lower than the preliminary values of Devine et al. (1984) for Tambora 
volatile degassing, but still place Tambora as the pre-eminent volcanic pollution event in historic 
time, with a total mass of 2.1x10'' kg of sulphur, chlorine and fluorine released to the 
atmosphere. Further studies of the poorly-constrained volume of the distal ash fall will probably 
lead to an increase in these estimates. In addition, we infer that about 2.8x10" kg of magmatic 
HO was introduced into the atmosphere during the eruption, or equivalent to more than doubling 
the stratospheric water-vapour content. Further addition of large quantities of meteoric water 
vapour to the stratosphere resulted from the large-scale convective flow of humid tropospheric 
air, entrained in the ascending eruption column. 

No measurements have been made of carbon dioxide levels in the Tambora products, but some 
inferences can be made of CO, output from the eruption. Magma of the type erupted from 
Tambora in 1815 is likely to have CO, levels of the order 500 ppm, judging from the solubility 
data of Stolper and Holloway (in press). Degassing of magma of this type would then yield about 
10% g CO, to the atmosphere during the 1815 eruption. Thus, the carbon dioxide output from 
Tambora would be roughly equivalent to the annual output from the Earth’s mantle, and only 
about 1% of the current annual anthropogenic output of CO . 

Sulphur Aerosol 

Sulphur output from Tambora during the three-day period in 1815 was more than double the 
current annual total sulphur output of volcanoes, which has been estimated as 0.9 to 1.2x10° 
kg/yr (Berresheim and Jaeschke 1983; Stoiber et al. 1987). In comparison, the annual global 
anthropogenic emission rate of sulphur dioxide is estimated as 1.3x10'' kg (Bach 1976). The fates 
and atmospheric effects of anthropogenic and volcano-derived sulphur aerosols are, however, 
quite different. The anthropogenic emission, caused by burning of fossil fuels, is mostly confined 
to the troposphere, where its residence time is short. In contrast, highly energetic explosive 
volcanic eruptions transport sulphur and other volatile species rapidly to the upper troposphere 
and lower stratosphere. In the case of Tambora, the early plinian events in April 1815 had 
sustained eruption columns of 33 to 43 km height above the volcano, but the convective columns 
during the main ignimbrite phase were probably in the 15 to 20 km range. With estimated magma 
source rate of 5.4x10* kg/s during the 10 April eruption, the source rate of volatiles to the 
atmosphere during the is period is calculated as 1.7x10° kg/s for sulphur, 4x10° kg/s for chlorine, 
2.7x10° kg/s for fluorine and about 10’ kg/s for magmatic water vapour. 

Sulphur emitted by Tambora was initially in the gaseous state, probably dominantly as SO, and 
lesser amounts of H,S and OCS, which are the precursor gases to sulphate aerosols and consume 
OH radicals. The large mass of magmatic and atmospheric water vapour injected into the 
stratosphere during the eruption (2.8x10'" kg) is a major potential source of the OH. Upon 
mixing with air, the sulphur dioxide would undergo oxidation to SO, and react with water vapour 
in the atmosphere to form an aerosol of sulphuric acid droplets. Reactions of the following type 
may account for the conversion of sulphur gases to sulphuric acid aerosol particles in the 


SO) OH > HOSO2 Oo-=50) HO, 
SOys# 11/2) OS: 

SO, + H,O > H,SO, (liq) 

HS’ +:3/2 0) => HLOnr SO, 

H,S + 20, > H,SO, (liq) 

The above mass estimates of volatile output from the eruption refer to elemental concentration 
of sulphur, chlorine and fluorine. Direct analysis of modern volcanic aerosols shows that they 
are typically composed of a 75% H,SO, aqueous solution (Hofmann and Rosen 1983). Converting 
the above petrologic estimate of 4.3x10'° kg elemental sulphur to sulphuric acid aerosol, we 
therefore estimate the Tambora sulphur-rich aerosol mass as 1.75x10"' kg, or an order of 
magnitude larger than the 1982 El Chich6n aerosol (McCormick and Swissler 1983). By 
comparison, Hammer et al. (1980) estimate a Tambora volcanic aerosol of 1.5x10"’ kg on the 
basis of the 1816 acidity layer in Greenland ice cores, and Stothers (1984) estimates 2x10" kg 
based on observed atmospheric effects. The difference in these estimates is within the 
uncertainties of the methods, but several factors make the petrologic estimate a minimum value. 
Firstly, further studies of the thickness and distribution of the distal tephra fall deposit preserved 
on the ocean floor may conceivably double the total erupted mass estimate and thus double the 
estimate of sulphur yield to the atmosphere. Secondly, the Tambora gas emission also involved 
about 1x10"! kg HC@ and 7.4x10"° kg HF, and the possible involvement of these gases in aerosol 
formation cannot be ruled out. Thirdly, the petrologic estimate is only of volatiles exsolved from 
the magma at the time of eruption, and does not include a possible separate volatile phase. 
Finally, the Tambora stratospheric aerosol or "dust cloud" also contained some particles of 
volcanic glass, as demonstrated by the recent identification in a South Pole ice core of Tambora 
glass fragments by microprobe analysis (J. Palais, personal communication). 

The Halogens 

The large-scale introduction of odd-chlorine species into the stratosphere during the 1815 
Tambora eruption is important because of the potential of chlorine in catalyzing the removal of 
O, and thus damaging the Earth’s ozone layer. That layer shields the biosphere from the effects 
of damaging solar ultraviolet radiation, such as effects on DNA and the immune-system response, 
skin cancer and sunburn. It is generally believed that diffusion of anthropogenic 
chlorofluoromethanes (CFC) from the troposphere is currently the principal source of 
stratospheric chlorine, but the importance of volcanic emissions as a potential source of 
stratospheric chlorine was first pointed out by Stolarski and Cicerone (1974). 

Stolarski and Butler (1978) estimated a stratospheric injection rate for volcanic chlorine of 
1.3x10’ kg/yr, or more than three orders of magnitude less than the 10'' kg HC emission during 
the 1815 eruption alone. By comparison, the annual release of chlorofluorocarbons is about 7x 10° 
kg/yr, and the budget of stratospheric chlorine is about 10° kg/yr. HC is generally the principal 
chlorine molecule in volcanic gases, but studies of the 1980 Mount St. Helens stratospheric cloud 
show that concentrations of methyl chloride (CH,C?) were as high or higher than concentrations 
of HC@ (Inn et al. 1981). HC2 is highly soluble in water, so possibly large quantities of the 
emitted HC? are dissolved in eruption-cloud water and returned to the surface of the Earth as 
precipitation during or shortly after eruption. 


Although large quantities of chlorine and fluorine are shown to be emitted by Tambora, it should 
not be assumed that these gases form aerosols in the stratosphere, as physical and chemical data 
indicate that HC? and HF gases are unlikely to form liquid aerosols under normal stratospheric 
conditions (Miller 1983; Solomon and Garcia 1984). As shown by Oskarsson (1980) (Figure 8), 
however, halogen aerosols may conceivably form at higher temperatures in the eruption column, 
and the presence of elevated concentrations of HC? and HF in volcanic-acidity layers in 
Greenland ice cores suggests that halogens have indeed become incorporated into some volcanic 
aerosols. Thus, the acidity layer from the 934 A.D. Eldgja eruption in a Greenland ice core 
contains at least 65% HC? (Hammer 1980). Herron (1982) has also shown high levels of both 
Cé and F in another Greenland ice-core layer from this eruption. Similarly, Herron (1982) and 
Hammer (1977) have both noted elevated C? levels in the Greenland ice-core acidity layer from 
the 1783 Laki eruption. Finally, very high C? concentration in a northwestern Greenland ice 
core, which was attributed by Herron (1982) to early nineteenth century volcanic activity, may 
conceivably represent material from the Tambora eruption. The ice-core data thus suggest that 
C2 and possibly F may enter the volcanic aerosol. This may not imply the formation of a discrete 
halogen aerosol, but rather that HC? and HF may be absorbed and dissolved in the sulphuric acid 

HC¢ is inert toward ozone, but reaction of HC? with OH leads to formation of atomic chlorine, 
followed by the catalytic decomposition of the ozone by the Cf. Thus in the stratosphere, C? can 
be released from HC? by reactions of the type: 

HO@41OHi= HO) Cr 

Similarly, methyl chloride can produce atomic chlorine by photolytic decomposition and attack 
by OH. Several reactions involving gaseous chlorine have the effect of converting odd-oxygen 
molecules (including ozone) to diatomic oxygen by C0 catalysis. They are reactions of the type: 


The only attempt to model the effects of large volcanic chlorine emission on the ozone layer was 
made by Stolarski and Butler (1978), who concluded that a Krakatau-size emission, involving 
3x10* kg C?, would result in about 7% depletion of the ozone layer. Chlorine output was two 
orders of magnitude higher than this value during the Tambora eruption, and major ozone 
depletion cannot be ruled out. Given the great importance of the ozone layer to the biosphere and 
climate, the modelling of the potential impact on atmospheric chemistry by a Tambora-size 
eruption is timely. 

Nothing is known about the possible atmospheric or environmental effects of the large (7.4x10"° 
kg) HF gas emission during the eruption indicated by our petrologic study. In general, HF is 
assumed to be very inert in the stratosphere. The photolysis of HF is shielded by oxygen, and 
the reaction of HF with OH is endothermic, so that it is believed that F atoms do not play the 
same role in stratospheric chemistry as chlorine atoms (Sze 1978). Furthermore, fluorine and to 
some extent chlorine, are known to adsorb onto tephra particles and thus may be rather rapidly 
removed from the atmosphere in the tephra fallout (Rose 1977; Oskarsson 1980). 










B fl 


eet b dbs dds 


Figure 8: Evolution of volcanic volatiles within an explosive eruption column, showing volcanic volatile 
reaction zones (from Oskarsson 1980). In the salt formation zone A, aerosol salt particles are 
formed at magmatic temperatures during high-temperature degassing of magma in the vent 
region. At temperatures in the range 338° to 700°C, surface adsorption of halogen gases 
occurs as they react with silicate material (adsorption zone B). At temperatures below 338°C 
sulphuric acid condenses as an aerosol in the condensation zone C. 

Fate of Volatiles in the Eruption Column and Atmosphere 

An explosive volcanic eruption represents a rapid transfer of heat and mass into the Earth’s 
atmosphere, resulting in a major thermal and chemical perturbation. In the case of the Tambora 
eruption, the thermal energy release alone was equivalent to about 1.3x10” ergs, most of which 
was introduced into the atmosphere over a period of about three or four days. Most of this energy 
was expended in convective mixing of the eruption column with ambient air and heating of the 
entrained air, resulting in the buoyant rise of the eruption column to heights of 43 km, as a 
mixture of pyroclastic fragments, volcanic gases and humid tropospheric air. 


Observations and theory (Sparks and Wilson 1982) shows that the solid particle weight fraction 
in high-eruption columns (1-n,) is only of the order 0.018; the remainder being almost entirely 
entrained atmospheric air and expanding volcanic gases. Assuming that most of the tephra that 
generated the fallout deposit (5.8x10"° kg) had entered the lower stratosphere, the mass of 
associated air lofted to the stratosphere would then be about equal, and equivalent to 
approximately 7x10"° m* at the surface. The water content of saturated air at 1 atm and 14°C is 
about 0.01 kg H,O/kg air. Thus the total mass of atmospherically-derived water entrained into 
the stratospheric eruption column could have been as high as 5x10"' kg. A portion would 
condense with rise in the eruption column and cause precipitation, but some would enter the 
stratosphere. Although large, this figure is only one-third of the mass of magmatic water 
introduced into the atmosphere (1.7x10"* kg), as discussed previously. Normally the content of 
water vapour decreases with height due to lowering of both temperature and saturation vapour 
pressure and condensation. However, water vapour is likely to be introduced to high levels under 
the conditions of elevated temperatures and turbulence within a buoyantly rising eruption column. 

Water vapour introduced to the stratosphere by an eruption column could be a major source of 
OH radicals by reaction of water vapour with photodissociated oxygen atoms. Evidence from 
ground-based spectroscopic measurements of OH during the 1982 El Chich6n eruption indicates 
that water vapour was injected at the level of 20 ppm (two to four times normal), and may have 
been responsible for the large ozone depletion observed in 1982-1983 (Burnett and Burnett 1984). 
Elevated levels of volcanically-derived OH from Tambora may have played a major role in 
generation of H,SO, by reaction with SO,, in the regeneration of free C? atoms from HC? and 
in direct reactions with stratospheric ozone. 

The field evidence indicates that during the main ignimbrite phase of the Tambora eruption, 
transport to the atmosphere was effected by two processes: the eruption column rising above the 
centre of the volcano, and secondary eruption columns rising from the coastline around the 
volcano as hot pyroclastic flows entered the ocean and flashed seawater to steam. About 35% of 
the erupted products entered the ocean in this manner and contributed to the secondary columns, 
probably resulting in a coastal ring of composite eruption columns around the entire Sanggar 
Peninsula, some 50 km in diameter. 

Evidently, extremely variable conditions exist in the eruption column, with great range in 
temperature, and mixing proportions of ambient air, condensed water vapour, volcanic gases and 
pyroclasts. Temperatures will range from magmatic (about 950°C) to stratospheric air (-60°C). 
The fate of the volcanic gases in the eruption column depends on temperature-dependent reactions 
in the atmosphere, and although conditions can clearly be highly variable, Oskarsson (1980) has 
recognized three zones (Figure 8). 

Salt Formation Zone 

A spontaneous non-equilibrium degassing occurs during a rapid pressure drop such as an 
explosive eruption. In the hottest core of the eruption column, within the jet-like mixture of 
pyroclasts and volcanic gases, aerosol salt particles are formed at near-magmatic temperatures. 
These are solids condensing from a magmatic gas phase and the solid reaction products of 
magmatic gas and its surroundings (Oskarsson 1980). Major sources for the salts are alkali 
metals, calcium, aluminum and silica from the silicate melt, and the reactive gases SO,, HC@, 
HF and NH,. The dominant products in the salt formation zone are chlorides, fluorides and 
sulphates of calcium and the alkali metals. Owing to the high vertical mass flow rates, particles 
formed in this zone are likely to represent a small fraction of total aerosol production, and they 


will be transported as suspended load high into the eruption column and downwind from the 

Surface Adsorption Zone 

As shown experimentally (Oskarsson 1980), the halogen gases react with silicate ash by surface 
adsorption and condensation of the gas phase at temperatures below 700°C. The reactions of the 
halogen gases with the glassy tephra will produce components such as calcium fluorosilicates, 
sodium and calcium chlorides and sodium fluoride. Halogens adsorbed on tephra particles in this 
zone will be removed relatively quickly from the eruption column during fallout. Thus, Rose 
(1977) has demonstrated that 17% of the C? released in the 1974 Fuego eruption was stripped 
from the eruption plume by adsorption onto tephra particles. 

Experiments and observations of the 1970 Hekla eruption show that a large fraction of the 
fluorine is stripped from the high-temperature region of the eruption column (338 to 700°C) by 
adsorption onto tephra and thus incorporated in the fallout deposit near source (Oskarsson 1980). 
Most of the Tambora fluorine emission may have been removed by this process, leading to 
fluorosis and thus accounting for the observed death of livestock. During the 1970 Hekla eruption 
in Iceland, fluorine-rich fallout led to poisoning of large numbers of livestock up to 200 km from 
the volcano (Thorarinsson and Sigvaldason 1972). The tephra fall from the eruption was 
unusually rich in adsorbed fluorine (up to 2,000 ppm). The concentration of the adsorbed fluorine 
in the fallout deposit was directly dependent on surface area of the tephra grains, and thus the 
concentration increased with: decreasing grain size. The total mass of fluorine deposited is 
estimated as 3x10’ kg, corresponding to 700 ppm of the total erupted mass from Hekla 
(Oskarsson 1980). 

Condensation Zone 

As temperature in the eruption column falls below 338°C, sulphuric acid can condense as an 
aerosol by a process controlled by the rate of oxidation of SO, by atmospheric oxygen and 
reaction with water vapour. Below 120°C the halogen acids condense and may form an aerosol 
prior to condensation of water. The sulphuric acid aerosol droplets can act as a medium in which 
other acid components, such as HF, HC? and water vapour can be dissolved. In the presence of 
tephra particles, a portion of the condensed aerosol can be stripped with fallout from the eruption 
plume by adsorption onto the silicate ash. Rose (1977) estimates that up to 33% of the sulphur 
released by the Fuego 1974 eruption was removed from the atmosphere in this manner. 

Effect of the Sulphuric Acid Aerosol on Climate 

Pollack et al. (1976) have shown that the optical properties of tephra are distinct from those of 
volcanic aerosols such as sulphuric acid, derived from conversion of volcanic gases. The 
importance of this was demonstrated during the 1980 Mount St. Helens eruption, when it was 
observed that the atmospheric cloud was composed dominantly of a sulphuric acid aerosol a few 
days after the eruption. During this eruption, the causes for the relatively short atmospheric 
residence time of even the finest-grained tephra were discovered to be due to silicate particle 
aggregation (Carey and Sigurdsson 1982). Because of these effects, apparently the potential 
climatic impact of a volcanic eruption is not primarily governed by the degree of explosivity or 
the volume of erupted magma, but more importantly by the chemical composition of the magma. 
Thus, recent studies indicate that the climatological effects of volcanic aerosol emission from 
large basaltic fissure eruptions may in fact be more important than the effects of explosive 
eruptions of silicic magmas (Sigurdsson 1982). 


It is generally accepted that the remarkable global meteorological and optical phenomena, 
observed months and years after the Tambora eruption, had a strong connection with activity of 
the volcano (Figure 9). Most of these phenomena can be attributed to the effect of the 
stratospheric volcanic aerosol. Owing to the sparse meteorological data available, the annual 
deviation of the global mean temperature due to the eruption is not well known, but spotty data 
indicate a minimum deviation in 1816 of -0.7°C in the northern hemisphere (Stothers 1984). In 
a reconstruction of long time series of temperature data from the eastern United States (Landsberg 
et al. 1968), the great climatic anomaly of the year 1816 is a unique event that also persists in 
1817 (Figure 10). Summer temperature was about 1.5°C below the 200-year average, and the 
June 1816 temperature about 3°C below average. 






0 I D 3 - 

Figure 9: Change in excess visual extinction (in astronomical magnitude units) following the 1815 
Tambora eruption at northern latitudes (after Stothers 1984). 




1810 1812 1814 1816 1818 1820 1822 

Figure 10: Observed climatic response following the Tambora 1815 eruption. Upper curve is annual 
summer temperature data for the eastern United States, at the latitude of Philadelphia, 
Pennsylvania, based on several long temperature series. The solid horizontal line shows the 
224-year average summer temperature (after Landsberg et al. 1968). Lower curve is annual 
June temperature data for New Haven, Connecticut. The lower horizontal solid line shows the 
145-year New Haven June mean temperature (World Weather Records 1927). 

Devine et al. (1984) and Palais and Sigurdsson (1989) evaluated the possible effect of volcanic 
eruptions on climate, and proposed a relationship between the mass yield of sulphur to the 
atmosphere from an eruption and the observed decrease in mean northern hemisphere surface 
temperature in the one to three years following the eruption, on basis of published temperature 
data (Figure 11). Palais and Sigurdsson (1989) found that the mean surface temperature decrease 
was related to the estimate of sulphur yield by a power function (r=0.92), with the power to 
which the sulphur mass is raised being equal to 0.308. Although these results appear to confirm 
a relationship between volcanic sulphur aerosol formation and climatic change, we emphasize that 
the temperature deviations are associated with large errors. 

As pointed out by Eddy (1988; this volume), the Tambora eruption was coincident with a 
depression in solar activity between about 1790-1830, i.e., the Dalton Minimum or the Little 
Maunder Minimum in sunspot numbers and aurorae (Figure 13). During these decades the 
characteristic 11-year cycle in solar activity persists, but the amplitude is reduced by an order of 
magnitude or more (Siscoe 1980). Variations in sunspot frequency have been linked to changes 
in the solar "constant", and in turn related to climatic changes. Thus the great reduction in 
surface temperature on Earth between 1650 and 1730 ("the Little Ice Age") corresponds to the 
Maunder Minimum, when there was a sudden reduction in sunspot numbers, almost to zero. It 
therefore appears likely that climate was already deteriorating by the beginning of the nineteenth 
century, due to reduction in solar activity. This climatic trend was then greatly amplified by the 
impact of the Tambora volcanic aerosol, culminating in the "year without summer" in 1816. Both 


AO" data on ice cores and northern hemisphere decadal temperature trends support the contention 
that a climatic change had set in by the first decade of the nineteenth century (Figure 12). Thus, 
for example, evidence from Peruvian ice cores shows that the decade 1810-20 is characterized 
by the most negative AO’ values (coldest temperatures) of the entire record (Figure 12), 
culminating in the southern hemisphere wet season of 1819-20 (Thompson et al. 1986; Thompson 
and Mosley-Thompson, this volume). The relative contribution of solar variability versus volcanic 
aerosol to the deterioration occurring after 1815 is unknown, but John A. Eddy (personal 
communication) estimates that solar variability may account for at most 10 to 50%. 


y oC =5.89e-5 * x SULFUR“0.308 R=0.92 



19019 10! 10!2 10'3 10'4 10'S 


Figure 11: The observed relationship between sulphur yield to the atmosphere during large volcanic 
eruptions and the northern hemisphere temperature decline following the event. Sulphur data 
are from Devine et al. (1984) and Palais and Sigurdsson (1989). Climatological data are from 
Rampino and Self (1982) and other sources cited in the text. The equation describes the best 
fit to the data, with a correlation coefficient of 0.92. 


05 Northern Hemisphere decadal temperature departures 
from the 1881-1975 mean 


Quelccaya summit core decadal oxygen isotope averages 
(1880-1980 mean) 1815 

1600 1700 1800 1900 
Year A.D. 

Figure 12: Variations in AO" in Peruvian ice cores and northern hemisphere surface temperature trends, 
showing surface temperature decline in progress before the onset of the Tambora 1815 eruption 
(after Thompson ef al. 1986). 


1780 1800 1820 1840 1860 1880 

Figure 13: Auroral and sunspot trends from 1780 to 1880, showing the "Dalton Minimum" or "Little 
Maunder Minimum" in solar activity between 1790 and 1830 (after Siscoe 1980). (A) Sunspot 
numbers and number of days per year on which aurorae were recorded in Norway. (B) Sunspot 
and auroral data in the United States and Europe, south of 54°N Latitude. 



The dynamics of the Tambora 1815 eruption columns and source rates of magma and volatiles 
can be determined by studying the deposits and petrology of the products. The initial plinian 
eruption of Tambora on 5 April was a brief but highly energetic event with eruption rate of 
1.1x10* kg/s producing a column height of 33 km. In the early phase of the paroxysmal eruption 
on 10 April, a plinian column rose to 43 km, with eruption rate of 2.8x10* kg/s. The buoyant 
column was only sustained for about three hours, before column collapse occurred due to 
increasing eruption rate. Subsequent ignimbrite-phase activity during a three-day period was at 
rates of about 5.4x10® kg/s, producing a co-ignimbrite fall deposit of 5.8x10"° kg and a 
pyroclastic flow deposit of 8.2x10" kg, or a total deposit of 1.4x10'* kg. The convective column 
above the volcano during the main ignimbrite phase was at least 20 km high, judging from grain- 
size data of the deposit, and thus injected material into the lower stratosphere. Although the ash 
fallout affected a broad area, the dispersal was dominantly to the west of Tambora, over Java and 
as far as Sumatra. This spread of the eruption plume is consistent with 10-year average rawin- 
sonde data for Surabaja in eastern Java, which shows dominant easterly upper troposphere and 
lower stratosphere winds for the spring months, with mean velocities ranging from 5 to 10 m/s. 
The column height evidence indicates that only about 2% of the erupted mass was emplaced into 
the middle stratosphere, up to 43 km, and that the vast majority of the erupted products were 
injected in the lower stratosphere and upper troposphere. 

The Tambora magma was enriched in volatile components, with 2 to 2.4 wt.% H,O, 570 ppm 
sulphur, 2,220 ppm chlorine and 570 ppm fluorine. Judging from the difference in volatile 
concentration in glass inclusions and in matrix glasses of the tephra, the yield of sulphur to the 
atmosphere was 4.3x10" kg, 10" kg of chlorine, and 7x10" kg fluorine. Magmatic water 
evolved from the volcano was about 2.8x10'* kg, whereas the mass of atmospheric water 
entrained in the eruption columns is estimated at 5x10'’ kg. Source rates of the volatile species 
were about 1.7x10° kg/s for sulphur, 4x10° kg/s for chlorine, 2.7x10° kg/s for fluorine and 10’ 
kg/s for magmatic water. 

Generation of the sulphuric acid aerosol by gas to particle conversion was probably greatly 
facilitated by OH radicals in the eruption cloud, derived dominantly from reactions between 
excited atomic oxygen and magmatic water vapour. Assuming a typical volcanic aerosol 
composed of 75% H,SO, and 25% water, the petrologic data indicate a minimum Tambora 
aerosol mass of 1.75x10"' kg. This compares closely to aerosol estimates based on the ice-core 
acidity layer (Hammer 1980) and atmospheric phenomena (Stothers 1984). 

The very high proportion of halogens released by the Tambora eruption is typical of volcanic 
activity of such trachytic magmas in subduction-zone environments. The fate of volcanic halogens 
in the atmosphere is unclear at this stage. Fluorine and chlorine most likely form HC? and HF 
gas molecules upon degassing from the magma. The latter is relatively inert in the stratosphere, 
as HF photolysis is shielded by oxygen and HF is also relatively indifferent to OH abundance. 

Chlorine was probably also removed in significant amounts from the high-temperature region of 
the Tambora eruption column by adsorption onto tephra. Studies of the 1974 Fuego eruption 
indicate that up to 17% of the chlorine was removed by this process (Rose 1977). Although HC? 
is not known to form stratospheric aerosols, chlorine may conceivably enter other aerosol 
droplets. Studies of ice cores cited above indicate that acidity layers from some eruptions contain 
significant chlorine, requiring incorporation of this species into the aerosol by some process. 
While HC? is relatively inert in the stratosphere, reaction with OH or by photolytic reactions 


leads to formation of atomic chlorine. As the Tambora eruption cloud was dominantly in the 
region below 30 km, which is photolytically inactive, formation of C? and ClO by the latter 
process would have been minor. On the other hand, we contend that water vapour was injected 
in large quantities, involving both magmatic and atmospheric water. Thus, OH radicals were 
abundant in the eruption column and available for reaction with HC? to produce atomic chlorine. 
Reactions of atomic chlorine with ozone are catalytic, and a single chlorine atom may destroy 
thousands of ozone molecules before it becomes inert and enters the HC? reservoir. Independent 
of their role in generation of single chlorine atoms, OH radicals from the eruption cloud would 
also lead directly to ozone destruction. 

The dominant environmental effects of the Tambora eruption were therefore probably of three 
types: (1) formation of a sulphuric acid aerosol, leading to a northern hemisphere temperature 
reduction of at least 0.7°C at the surface and stratospheric heating; (2) adsorption of fluorine onto 
tephra, leading to very high fluorine levels in the fallout on the ground in Indonesia and resulting 
in widespread fluorosis; and (3) extensive ozone depletion as a consequence of generation of odd 
chlorine atoms and high levels of volcanically-derived stratospheric OH radicals. 


This research was carried out with funding from the National Science Foundation (grants EAR- 
8607336 and EAR-8804117), and field studies in Indonesia were made possible by funding from 
the National Geographic Society (grant NGS 3390-86). We thank the Volcanological Survey of 
Indonesia for collaboration in the field, and the Indonesian Research Council (LIPI), for 
permission to undertake research in Sumbawa. The assistance of David Browning in electron 
microprobe and grain-size analysis is gratefully acknowledged. 


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The Possible Effects of the Tambora Eruption in 1815 on Atmospheric 
Thermal and Chemical Structure and Surface Climate 

R.K.R. Vupputuri’ 


A coupled 1-D radiative-convective-photochemical diffusion model that takes into account the 
influence of ocean inertia on global radiative perturbation is used to investigate the possible 
climatic and other atmospheric effects of the Tambora eruption of 1815. The volcanic cloud was . 
introduced in the model stratosphere between 20-25 km, and the global average peak aerosol 
optical thickness was assumed to be 0.25. Both the aerosol optical thickness and aerosol 
composition determining the optical properties were allowed to vary in the model atmosphere 
during the life cycle of the volcanic cloud. The results indicate that the global average surface- 
temperature decreases steadily from the date of eruption in 1815 with maximum cooling of 1°K 
occurring in spring 1816. The calculations also show significant warming of the stratosphere, 
with temperature increasing up to 15°K at 25 km in less than six months after the date of 
eruption. The important effects of the Tambora eruption on stratospheric ozone and UV-B 
radiation at the surface are also discussed. 


Modelling studies of global radiative perturbations caused by volcanic eruptions are extremely 
important to understand the nature of past and present changes in the Earth’s climate. As pointed 
out by Kondratyev (1983), the primary mechanism by which volcanic activity influences the net 
radiation (and consequently the climatic system) is through alterations in the aerosol content of 
the atmosphere. It is well known that volcanic aerosols scatter and absorb solar radiation and 
absorb and emit infrared radiation. The net effect causes general cooling of the troposphere and 
the surface, and warming in the stratosphere. The extent of the cooling and warming, however, 
depends upon the composition, size distribution and the morphological structure of aerosols. 

Several prominent volcanic eruptions took place during the past 200 years (Laki, 1783; Krakatau, 
1883; Mount Agung, 1963; and El Chichén, 1982). The largest and deadliest was that of Mount 
Tambora in April 1815, on the island of Sumbawa, Indonesia (8°S, 118°E). It was also the 
world’s greatest ash eruption since the end of the last ice age. The dust veil index (a measure of 
increase in the atmospheric turbidity arising from small particles injected into the stratosphere) 
has been estimated to be more than twice that of Agung (Lamb 1970; Robock 1981, Mitchell 
1982). The Tambora eruption is blamed by some studies in the literature for the cold summer of 
1816 on the east coast of North America, where average temperature was the lowest on record 
with 1.5 to 2.5°C below the seasonal norm (Landsberg and Albert 1974). Indeed 1816 was called 
"the year without a summer" in New England and eastern Canada, where daily minimum 
temperatures were abnormally low from late spring through early autumn. It was also very cold 
and wet in western Europe in the summer of 1816, although it was milder at some stations in 
eastern Europe. Despite these earlier claims of strong cooling on a regional basis, the more recent 
analysis of climatic data for 1781-1983 by Angell and Korshover (1985) suggests that there is no 

' Canadian Climate Centre, 4905 Dufferin Street, Downsview, Ontario M3H 5T4, Canada. 


clear evidence of strong cooling on a hemispheric basis following the Tambora eruption. As 
pointed out by Angell (1988), the reason for the lack of evidence of strong volcanically-induced 
cooling on a hemispheric basis is that such cooling may or may not have been observed 
depending upon the extent of sea-surface warming in the eastern equatorial Pacific due to an El 
Nifio event following the volcanic eruption. 

It is clear from the above arguments that the injection of ash, sulphur and dust into the 
stratosphere by a large volcano could alter the existing radiative-photochemical balance of the 
Earth’s atmosphere, in turn leading to changes in the vertical temperature and chemical structure 
and surface climate. In this respect a volcanic event as large as the Tambora eruption provides 
a unique opportunity for a case study of the response of the climatic system to a global radiative- 
photochemical perturbation. It also allows testing of our ability to model and understand the 
nature of the climatic system and climatic change. Several model calculations have been made 
in the past to study the climatic impact of Agung and El Chichén eruptions (Hansen et al. 1977; 
Robock 1984; McCracken and Luther 1984; Vupputuri and Blanchet 1984). All these calculations 
showed warming of the stratosphere and cooling in the troposphere and at the surface, although 
the amplitudes of warming and cooling differ depending upon the assumed peak aerosol optical 
depth, altitude of peak aerosol concentration and optical properties. In this paper a 1-D time 
dependent radiative-convective-photochemical diffusion model (RCPD model) taking into account 
the thermal inertia of oceans is used to investigate the thermal and chemical response of the 
atmosphere and surface climate to radiative-photochemical perturbations caused by the Tambora 
eruption in 1815. 

The Climatic Model 

The coupled one-dimensional climatic model extending from the surface to 60.5 km is described 
in detail in Vupputuri (1985). It involves combining the radiative-convective model of the type 
developed by Manabe and Wetherald (1967) with a photochemical transport model. Starting from 
the assumed vertical temperature distribution and chemical composition, the basic procedure is 
to compute the local net radiative heating and cooling and photochemical sources and sinks at 
each altitude to determine the vertical temperature and trace-constituent structure with the time 
marching method. The upward heat transfer by atmospheric motions is taken into account 
implicitly by a simple numerical procedure called convective adjustment - first introduced by 
Manabe and Strickler (1964). Using this numerical procedure, the vertical lapse rate is restored 
to a pre-assigned stable value (6.5° km") whenever it becomes greater due to radiative heat 
transfer. The relative humidity is fixed in the model, and the mixing ratio of water vapour is 
computed as a function temperature. Cloud-top altitude is assumed to be at 6.5 km, with a 
fractional cloud cover kept at 50%. The climatic model is coupled to the underlying surface 
through the energy balance equation at the surface. For climatic change calculations in this study, 
thermal heat capacity of the underlying surface is assumed to zero for the land and the value 
appropriate for the upper mixed layer in the case of the ocean. For several other computational 
details of the model, see Vupputuri (1985). 

The Radiative Transfer Model 
The solar radiation code used to compute the short wave solar heating within the atmosphere is 
based on the delta-Eddington method, which is computationally efficient and fairly accurate 

(Joseph et al. 1976). It considers the absorption and scattering by atmospheric gases (H,O, CO,, 
O, and NO,), aerosols and cloud droplets. To compute the infrared cooling due to H,O, O, and 


CO,, the analytical formulae for the mean transmissivities of finite frequency intervals derived 
by Kuo (1977) have been adopted. The mean transmission functions take into account the 
temperature effect and overlapping absorption between gases, and the computed transmissivities 
have been shown to be in good agreement with line-by-line calculations. For other trace gases, 
such as N,O, CH, and CFCs, empirical expressions for the mean band absorptivities based on 
laboratory and spectroscopic data (Burch et al. 1962; Ramanathan et al. 1985) were adopted. 
Also considered in long-wave calculations are the heating or cooling-rate contributions due to 
aerosols in the atmospheric window band. Both IR and solar radiation codes have been validated 
by comparing them with other standard radiation codes through participation in the workshop on 
the intercomparison of radiative codes in the climatic models (World Meteorological Organization 

The Photochemical Model 

The photochemical system considered here includes the important reactions affecting the 
concentration of ozone and other relevant trace constituents in the atmosphere above 10 km: they 
are listed in Table 1. The chemical species considered are: O, (O, O('D), O,), HO, (H, HO, 
HO,), NON, NO, NO,, HNO,), CH, NO, CF,C?f,, CFC?, and C?;(C?2, C20,.CENG)) 
chemical species. The concentrations of HO, CH, and H, are specified based on observations. 
The chemical kinetics and photochemical data used are based on NASA (1985) recommendations. 
Computed photodissociation rates take into account the effects of Rayleigh scattering and 
absorption and scattering by aerosols and cloud droplets. The chemistry of the global troposphere 
below 10 km is considered to be much more complex, and therefore to have more uncertainties 
than in the stratosphere due to the presence of higher hydrocarbons, heterogeneous processes and 
long photochemical relaxation times for the chemical species. In view of some of these 
uncertainties, the chemistry is frozen in the troposphere by prescribing the ozone concentration 
below 10 km for the purpose of this study. 

The Perturbed Aerosol Model for the Tambora Eruption 

To calculate possible climatic and other atmospheric effects of the Tambora eruption, 
observational information on perturbed aerosol concentration and optical properties as a function 
of time starting from the date of eruption are needed. It is not, however, possible to obtain such 
detailed information for Mount Tambora. For this study it is assumed that the dust veil index for 
Mount Tambora is roughly twice that of Agung. Using the peak optical depth of 0.125 as a 
representative global average value for the added aerosols in the case of Agung (Hansen et al. 
1978), the maximum optical thickness for Mount Tambora is estimated to be 0.25. The assumed 
shape of the vertical aerosol profile producing this maximum optical thickness is shown in 
Figure |. The vertical distribution of perturbed aerosols and the altitude of peak aerosol 
concentration are similar to those of Agung and El Chich6n eruptions. Since the added aerosol 
concentration and the optical properties are expected to vary with time during the lifetime of the 
volcanic cloud, it is not realistic to assume fixed optical thickness and properties for the model 
calculations. In the present calculations, both the aerosol optical thickness and properties are 
varied as a function time starting from the date of volcanic eruption. The assumed variation of 
perturbed aerosol optical thickness is illustrated in Figure 2. For the first four months after the 
eruption, the perturbation optical thickness is allowed to increase linearly to a maximum 0.25, 
and during this period ash optical properties are assigned for the added aerosols. From four to 


Table 1: The Principle Chemical and Photochemical Reactions Used in the Model. 

O,+hyv>O0+0 OH + O, > HO, + O, 
O0+0,+M>0,+M OH + O>H +0, 
O0+0,7>0, + O, HO, + O, > OH + 20, 
O,+hy>0,+0 H + O, + M>HO, + M 

O, + hy >O, + O('D) H + O,> OH + O, 

O('D) + 0, > 0, + O, HO, + HO, > H,O, + O, 
o(’’D) +M>O+M H,O, + hy OH + OH 

NO, + hy > NO +0 H,O, + OH > H,O + HO, 
NO + 0 +M->NO,+M OH + HO, > H,O + O, 

NO, + O>NO + O, HO + NO, + M>HNO, + M 
NO + O, > NO, + O, OH + HNO, -H,O + NO, 
N,O + hy>N, +0 HO, + NO > OH + NO, 
N,O + O('D) > N, + O, CF,C?, + hy > CF,Ce + Ce 
N,O + O('D) > NO + NO CFCf, + hy > CFC2, + Ce 
NO + hy>N+0O Gy Se lsh a lel es lol 
N+0O,>NO+0O Ce + CH, > HCé + CH, 

N + NO>N, +0 Ce + HO, > HCE + O, 

H,O + O('D) > OH + OH OH + HC? > H,O + Ce 

HO > hy > H OH HCé + OOH + Ce 

CH, + O('D) > OH + CH, HCé + hy >H + Ce 

CH, + O, + M>CH,O, + M Cf + O, > Clo + O, 

CH,0, + NO>CH,O + NO, Clo + O>Cf + O, 

CH,O + O,>H,CO + HO, Clo + NO>C¢ + NO, 
ILCO + hy > Ho + HCO Cfo + NO, + M>CéNO, + M 
HCO + 0,>CO + HO, Cé£NO, + hy > ClO + NO, 
OH + CO>CO, + H Cé£NO, + HC? > Ce, + HNO, 
H, + O('D) > H + OH Cf£NO, + O> CLO + NO, 
HNO, + hy > OH + NO, 

10 months the peak optical thickness remains the same but the optical properties are changed 
from ash to sulphuric acid. After 10 months the optical thickness decreases exponentially until 
it reaches the background stratospheric value while the optical properties change from sulphuric 
acid to background stratospheric aerosols. The optical parameters (extinction coefficients, single 
scatter albedo, asymmetry factors) for ash, sulphuric acid and background stratospheric aerosols 
vary as a function of wavelength both in solar and infrared spectra. Sulphuric acid properties 
chosen for this study are those reported in Bundeen and Fraser (1982), and they are derived 
assuming the aerosol particles are composed of 75% H,SO, and 25% H,O. The ash optical 
properties are determined by assuming an imaginary refractive index of 0.002 (Patterson and 
Pollard 1983). Both the sulphuric acid and ash properties are similar to those used for the El 
Chich6én volcanic eruption. 

Results and Discussion 
Before discussing the results of atmospheric response to radiative-photochemical perturbations due 

to the Tambora eruption, it should be pointed out that for the prescribed annual average insolation 
and background stratospheric aerosols, the coupled 1-D model produced reference atmosphere 


simulations of minor trace constituents and temperature that are representative of natural 
background atmosphere in tropical latitudes. A detailed comparison of reference atmosphere 
model simulations of ozone and temperature, and discussion on some of the deficiencies in 1-D 
model calculations, were given in Vupputuri (1985). 




14.53.19 THUR 23 JUN, 1988 

) 500 1000 1500 2000 2500 3000 

OT 1 

Figure 1: Assumed aerosol concentration profile (NO/CC) for Mount Tambora volcanic eruption which 
produces the maximum optical depth of 0.25. 

Effects on Solar and Infrared Radiation 

Figure 3 shows the calculated change in direct, diffuse and net solar radiation at the surface as 
a function of time beginning from the date of eruption of Tambora in 1815, while the 
corresponding effects on infrared flux and planetary albedo are illustrated in Figure 4. Direct 
solar flux decreases by about 15% following the eruption (Figure 3). However this decrease is 
compensated in large measure by an increase in diffuse radiation, leaving a net decrease of solar 
flux at the surface of about 4%. There are no observational data on direct solar radiation 
following the Tambora eruption. The visual extinction curve constructed by Stothers (1984) 
suggests that the excess zenithal visual extinction increases rapidly during the first four to five 
months from the eruption date, and then returns to normal gradually within four to five years. 
The time variation of visual extinction is quite consistent with the variation of calculated direct 
solar radiation following the eruption of Tambora. As indicated in Figure 4, the infrared flux at 
the surface also decreases by up to 4%, while planetary albedo increases by about 7% following 
the Tambora eruption. 



$ 0.20 
fy ae 
os fo 

3 =! . 
3 z 0.15 
2 O 
Pe ee 
se O 
8 0.10 
bed ae 

0.00 t pone 7 

= 1815 1816 1817 1818 1819 1820 

Figure 2: The assumed variation of perturbed stratospheric aerosol optical thickness with time following 
the Mount Tambora eruption in 1815. 


a i ee we ee DIRECT FLUX 
Se 2s ee Se Ba DIFFUSE FLUX 






db MnT, lyoy 


sU.U%eek PKL 

1815 1816 1817 1818 


Figure 3: The calculated change in direct, diffuse and net solar radion (in %) at the surface following the 
Mount Tambora eruption. 





3 7 0 i %y, seeceeeesee 


11.19.25 FRI 6 MAY, 1988 

T T T T Be Ne S| 
1815 1816 1817 1818 1819 1820 

OT | 

Figure 4: The calculated percentage change in the infrared flux at the surface and planetary albedo 
following the Mount Tambora eruption. 

Effects on Global Climate 

Figure 5 shows the calculated response of stratospheric and surface temperature as a function of 
time following the Tambora eruption. The response is shown for two different assumed heat 
capacities for the underlying lower-boundary surface. The solid curve represents the calculated 
surface-temperature response assuming that the lower boundary has no heat capacity (valid 
assumption for the land surface). The dashed curve, on the other hand, corresponds to the heat 
capacity of the underlying surface equivalent to that of 70 m of ocean water. The combined 
response of land and ocean is given by the dotted curve. In the stratosphere, the ocean heat 
capacity has no effect on temperature response due to volcanic forcing caused by the Tambora 
eruption (dash-dot curve in Figure 3). Note however that the stratosphere responds much more 
quickly to the global radiative perturbation due to volcanic forcing. As indicated in Figure 5, the 
eruption of Tambora could have resulted in a peak warming of about 15°K within less than six 
months after the eruption. But in the case of the troposphere and the surface it takes almost twice 
as long to reach the maximum cooling. The calculated maximum cooling for the land surface 
following the Tambora eruption is about 2°K, and for the combined land and ocean it is roughly 
1°K. It may be seen from Figure 5 that not only the land surface cools faster than the ocean 
surface but it also warms faster than ocean as the temperature recovers to its pre-volcanic state. 





SO Oe 
2 OT ea 




Figure 5: The calculated stratospheric and surface-temperature change (O K) following the Tambora 
eruption. The surface-temperature response is shown for two different assumed heat capacities 
for the underlying lower boundary surface. 

The physical explanation for the calculated temperature effects in the present model is 
Straightforward. For the aerosol concentrations and optical properties assumed in this study, the 
net effect of the added aerosols in the stratosphere from the Tambora eruption would be to 
increase the planetary albedo and decrease both the solar and thermal radiation at the surface 
(Figures 3, 4). This leads to cooling in the troposphere and at the surface. The local heating in 
the stratosphere, on the other hand, is caused by both the absorption of thermal radiation 
emanating from the warmer lower atmosphere and in situ absorption of solar radiation (in the 
near infrared and UV part of the spectrum) by the added aerosols. Due to low air density in the 
stratosphere, only a small change in radiational energy is needed to cause a large change in local 
air temperature. 

As pointed out earlier, although surface temperatures were abnormally low in the summer of 
1816 (following the Tambora eruption) in New England, eastern Canada and parts of western 
Europe, the evidence for a strong land-surface temperature cooling on a global average basis (as 
indicated by the results of this study) following the Tambora eruption is rather weak. As Angell 
and Korshover (1985) indicate, the extent of surface cooling on an hemispheric or global basis 
depends critically upon the timing and strength of El Nifio in relation to the time of a large 
volcanic eruption. Indeed Quinn et al. (1978) have found some evidence for a moderate El Nifio 
one year after the Tambora eruption. This might partially explain the reason for the lack of 
evidence of strong cooling on a global or hemispheric basis in 1816 following the Tambora 


Effects on Stratospheric Ozone and UV-B Radiation at the Surface 

As mentioned earlier, the absorption of solar and thermal radiation by the added aerosols in the 
stratosphere from a large volcanic eruption such as Tambora can lead to a large increase in the 
stratospheric temperature. The altered temperature in turn effects the concentration of ozone and 
other minor constituents through temperature-dependent reaction-rate coefficients. Due to the 
inverse relationship between ozone and temperature in the middle stratosphere, the temperature 
increase in that region causes ozone concentration to decrease. Ozone is also destroyed in the 
stratosphere by enhanced photodissociation resulting from the backscattered UV radiation from 
the added aerosols. Figure 6 shows the computed ozone column reduction and UV-B radiation 
increase at the surface following the Tambora eruption. It is seen from Figure 6 that the added 
aerosols in the stratosphere from the Tambora eruption could have resulted in up to about 7% 
decrease in total ozone, which translates into up to 15% increase in the UV-B radiation at the 
surface following the eruption event. Although there were no observational data following the 
Tambora eruption to verify the computed ozone depletion, the analysis of Umkehr observations 
following El Chichén in 1982 by DeLuisi et al. (1984) clearly indicates the evidence of 
volcanically-induced ozone depletion. This lends support for the theoretical calculations shown 
in Figure 6. 



qe : {fie Ee (Ca are SURFACE UV-B RADIATION 


1815 1816 1817 1818 1819 1820 

Figure 6: The calculated changes in the total ozone column and UV-B radiation at the surface (in %) 
following the Mount Tambora eruption. 


Concluding Remarks 

The Tambora eruption in 1815 - the largest and deadliest eruption in recorded history - also 
injected the greatest amount of ash, sulphur and dust into the stratosphere. The event that 
produced the largest dust veil index provided a unique opportunity to investigate climatic and 
other atmospheric response to global radiative perturbation, and to understand the effects of 
volcanic eruptions on past and present climate. 

Despite the simplicity of a 1-D radiative-convective-photochemical diffusion model that does not 
include the interaction of radiative heating perturbation with atmospheric dynamics and other 
uncertainties regarding the input data, the magnitude of land-surface temperature decrease 
generally agrees with the observed cooling in the east coast of North America, where 1816 was 
called "the year without a summer". However there is lack of strong evidence from the 
observational data to support the computed combined land-ocean surface cooling on hemispheric 
or global bases. The computed cooling may or may not have been observed depending upon the 
timing and extent of the El Nifio event following the Tambora eruption. There were no 
observational data in 1816 to verify the computed changes in ozone and temperature in the 
stratosphere following this eruption. However, the well-documented observational evidence of 
temperature warming (Quiroz 1983) and ozone depletion (DeLuisi et al. 1984) in the stratosphere 
following the El Chichén eruption lends support for the calculated changes in the case of the 
Tambora eruption. 

One should exercise caution in interpreting either the observations or the model results presented 
here. Temperature and ozone observations are not as detailed as desired for accurate 
determination of observed climatic and total ozone changes on a global average basis. Further, 
the observed climatic change is not a part of other natural events such El Niftio, QBO and the sun- 
spot cycle, or simply noise in the climatic system (Hansen et al. 1978). Climatic calculations in 
the model are too simplified and, in particular, the model is not capable of handling the complex 
interactions between radiative heating and large-scale dynamics and other cloud feedback 
mechanisms. Nevertheless, the calculated amplitudes of climatic and ozone perturbations resulting 
from Tambora’s eruption are large enough to believe that volcanic eruptions do indeed strongly 
affect the Earth’s climate and the ozone layer. As pointed out by Angell (1988), the reason that 
the evidence for volcanically-induced cooling of the Earth’s surface in the past was so uncertain 
is that the cooling may or may not have been observed depending upon the extent of warming 
due to El Nifio following the volcanic eruption. Detailed observations and careful data analysis 
taking into account the impact of other natural events following a major eruption such as 
Tambora would enable us to understand better the role of volcanic aerosols in altering the 
radiative-photochemical balance of the global atmosphere and climate. 


The author thanks: Dr. G.J. Boer and the Director General of the Canadian Climate Centre for 
encouragement and support; Mr. Frank Szeckeli and Lynda Smith for programming and 
manuscript preparation support, respectively. 


Angell, J.K. 1988. Impact of El Nifio on the delineation of tropospheric cooling due to volcanic 
eruptions..Journal of Geophysical Research 93:3696-3704. 


Angell, J.K. and J. Korshover. 1985. Surface-temperature changes following the six major 
volcanic episodes between 1780 and 1980. Journal of Climate and Applied Meteorology 

Bundeen, W.R. and R.S. Fraser. (eds.) 1983. Radiative effects of the El Chichén volcanic 
eruption: preliminary results concerning remote sensing. NASA Technical Memorandum 
84959, Goddard Space Flight Center, Greenbelt, Maryland. 

Burch D.E., D. Grynak, E.B. Singleton, W.L. France and D. Williams. 1962. Infrared 
absorption by CO,, H,O and minor atmospheric constituents. AFCRL-62-688, Ohio State 
University Contract AF19(604)-2366. 

DeLuisi, J.J., C.L. Mateer and W.D. Komhyr. 1985. Effects of the El Chich6n stratospheric 
aerosol cloud on Umkehr measurements at Mauna Loa, Hawaii. Atmospheric ozone. 
C.S. Zerefos and A. Ghazi. (eds.). D. Reidel Publishing Co., Dordrecht. 

Hansen, J.E., W.C. Wang and A.A. Lacis. 1978. Agung eruption provides test of a global 
climate perturbation. Science 199:1065-1068. 

Joseph, J.H., W.J. Wiscombe and J.A. Weinman. 1976. The delta-Eddington approximation for 
radiative flux transfer. Journal of Atmospheric Sciences 33:2452-2459. 

Kondratyev, K. Ya. 1983. Volcanoes and Climate. R.D. Bojkov and B.W. Boville (eds.). World 
Meteorological Organization, WCP-54. 

Kuo, H.L. 1977. Analytic infrared transmissivities of the atmosphere. Beitrage zur Physik der 
Atmosphaere 50:331-349. 

Lamb, H.H. 1970. Volcanic dust in the atmosphere, with a chronology and assessment of its 
meteorological significance. Philosophical Transactions of the Royal Society of London 


Landsberg, H.E. and J.M. Albert. 1974. The summer of 1816 and volcanism. Weatherwise 

Manabe, S. and R.F. Strickler. 1964. Thermal equilibrium of the atmosphere with a convective 
adjustment. Journal of Atmospheric Sciences 21:361-385. 

Manabe, S. and R.T. Wetherald. 1967. Thermal equilibrium of the atmosphere with a given 
distribution of relative humidity. Journal of Atmospheric Sciences 24:241-259. 

McCracken, M.C. and F.M. Luther. 1984. Preliminary estimates of the radiative and climatic 
effects of the El Chichén eruption. Geofisica Internacional 23(3):385-401. 

Mitchell, J.M., Jr. 1982. El Chichén, weather-maker of the century. Weatherwise 35:252-259. 
NASA Panel for Data Evaluation. 1985. Chemical kinetics and photochemical data for use in 

stratospheric modelling. California Institute of Technology, Jet Propulsion Laboratory, 
Publication 85-37, Pasadena, California. 217 pp. 


Patterson, E.M. and C.O. Pollard. 1983. Optical properties of the ash from El Chichén volcano. 
Geophysical Research Letters 10:317-320. 

Quinn, W.H., D.O. Zopf, K.S. Short and R.T.W. Kuo Yank. 1978. Historical trends and 
statistics of Southern Oscillation, El Nifo and Indonesian droughts. Fishery Bulletin 

Quiroz, R.S. 1983. The isolation of stratospheric temperature change due to the El Chichén 
volcanic eruption from non-volcanic signals. Journal of Geophysical Research 88:6773- 

Ramanathan, V., H.B. Singh, R.J. Cicerone and J.T. Kiehl. 1985. Trace gas trends and their 
potential role in climate change. Journal of Geophysical Research 90:5547-5566. 

Robock, A. 1981. A latitudinally dependent volcanic dust veil index and its effect on climate 
simulations. Journal of Volcanology and Geothermal Research 11:67-80. 

Stothers, R.B. 1984. The great Tambora eruption in 1815 and its aftermath. Science 224:1191- 

Vupputuri, R.K.R. 1985. The effect of ozone photochemistry on atmospheric and surface 
temperature changes due to increased CO,, N,O, CH, and volcanic aerosols in the 
atmosphere. Atmosphere-Ocean 23:359-374. 

Vupputuri, R.K.R. and J.-P. Blanchet. 1984. The possible effects of El Chichén eruption on 
atmospheric thermal and chemical structure and surface climate. Geofisica Internacional 

World Meteorological Organization. 1984. The intercomparison of radiation codes in climate 
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Meteorological Organization, WCP-93. 


Climatic Effects of the 1783 Laki Eruption 

Charles A. Wood! 


From 8 June 1783 to 7 February 1784, 12 km’ of lava poured from a series of volcanic vents in 
southern Iceland, devastating farmland and ultimately causing a severe famine that decimated the 
island’s human and animal populations. This Laki eruption apparently had much more widespread 
consequences, however, for its ash and sulphurous gases were transported in the lower 
atmosphere across Europe causing a remarkably warm summer, which was followed in Europe, — 
eastern North America and at least some parts of Asia by one of the most severe winters on 
record. Poor weather continued through the summer and winter of 1784. Scarce meteorological 
measurements and abundant written records and proxy data graphically document these climatic 

Conventional volcano-climate theories cannot readily explain these apparent climatic effects. Great 
eruptions such as Tambora, 1815 or Krakatau, 1883 explosively emplace volcanic aerosols into 
the stratosphere where, during a two- to three-year period before they are finally flushed out, 
they absorb incoming radiation, thus depriving the lower atmosphere of a portion of its heat. For 
the Laki eruption there is no direct evidence that significant quantities of sulphuric aerosols 
reached the stratosphere. If the continuing climatic deterioration of 1784 was related to the clearly 
volcanic weather of 1783, then a new mechanism needs to be identified. 


The first recognition that volcanism may effect climate was Benjamin Franklin’s (1784) 
speculation that the non-explosive Laki eruption in Iceland could be responsible for the poor 
climate in Europe and North America during the summer and winter of 1783. Although the Laki 
eruption was the largest effusive volcanic activity in historic times, it, and its possible climatic 
effects, have been studied only at the reconnaissance level. Nonetheless, general information on 
the nature and chronology of the eruption, and compilation of reports of anomalous weather 
thereafter, provide strong support for Franklin’s prescient theory, and lead to puzzling aspects 
of conventional eruption-climate relationships. 

Laki Eruption of 1783 

The only accessible accounts of the Laki eruption are Thorarinsson’s (1969) 20-year old 
preliminary report and a recent abstract by Thordarsson et al. (1987), which incorporate 
eyewitness descriptions and reconnaissance geological mapping. Most of the 1783 activity 
occurred along a 25-km-long series of fissures, creating large fields of lava flows with small 
cones marking their vents. The eruption began on 8 June 1783, following three weeks of 
earthquakes, and continued for eight months (until 7 February 1784). Thorarinsson estimates that 
~ 12 km’ of tholeiitic lava flows were emplaced, with the majority (10 km’) produced during the 
first 50 days (8 June to 28 July 1783). From examination of Thorarinsson’s map (Figure 1) 

' NASA Johnson Space Center, SN2, Houston, Texas 77058, U.S.A. 


showing the extent of the lava at different dates, it appears that a significant portion of the flows 
had formed by 21 June and the majority of the 10 km’ by 22 July. 

P 7 
AND " Ne 
1783 = 11 20f 




__- 900 



i) ¢ 7 600 
2 = ? EN 
s (0) ) ey oS ( Miho SA SN 
€ I | ee a VARMARFELL 
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: & 
Cw7* ( 

~ J 

ee WSIS ad iN 

—S o—— Vo 

pees) hex 
2 2) 2 5 aac 
? y ae! >} 

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Figure 1: Map of Laki from Thorarinsson (1969). The Laki lavas of 1783 are lightly stippled. 


Thus, the extrusion rate during the first two weeks may have been considerably greater than the 
50-day average of 2,200 m’*/sec estimated by Thorarinsson. Following this initial, high-rate 
extrusion, activity continued at a much slower rate for the next six months. The total area 
covered by the flows is 565 km? (Thoroddsen 1925), which implies an average thickness of ~ 20 

A recent development has been the recognition that the Laki fissure activity was part of volcanic- 
tectonic eruption centred on the Grimsvotn caldera, northeast of Laki (Thordarsson et al. 1987; 
Thordarsson and Self 1988). During and after the Laki fissure eruptions, Grimsvotn had a series 
of explosive eruptions between 18 July 1783 and 26 May 1785. 

From the perspective of climatic effects it is important to know how much explosive activity ~ 
occurred at the beginning of the eruption. According to eyewitnesses the eruption was very 
violent during the first few days, with enormous lava fountains. Thorarinsson pointed out that 
groundwater may have been involved in some of these eruptions because phreatomagmatic tephra 
cones were formed. Based upon his field measurements of buried ash from Laki, Thorarinsson 
believed the volume of explosively erupted material was 0.3 km’; he discounted earlier estimates 
of 3 km*. Thordarsson et al. give a similar value (0.21 km’) for the tephra. 

Effects of the Laki Eruption in Iceland 

Thorarinsson (1969) states that the Laki eruption was the greatest catastrophe in Icelandic history, 
and the mortality statistics bear him out. Gases from the eruption stunted the growth of grass so 
that it was insufficient to feed livestock. As a result, 50% of the cattle, 79% of the sheep, and 
76% of the horses starved. During the next three years 24% (9,000 people) of the human 
population died of starvation, and the population did not return to earlier levels until 1780 
(Jackson 1982). 

Ogilvie (1986) cites contemporary diaries that provide graphic information on the effects of the 
eruption. From 8 June to at least 26 August, "the air was full of ash and smoke. On the rare 
occasions that we have had a glimpse of the sun it has looked like the reddest blood". Grass 
turned yellow and white due to "sulphurous rain" and it withered to the roots. Fishermen were 
not able to go to sea because of "murmurings" (earthquakes?) and continuous smoke that reduced 
visibility to less than a mile. In northeastern Iceland, one writer recorded, "From early June, and 
to this time (13 August) we have lived in continual smoke and fog, sometimes accompanied by 
sulphur-steam and ashfalls." 

Weather in Iceland during the eruption seemed variable from location to location according to 
written records analyzed by Ogilvie (1986), but she concludes that 1783-84 winter began very 
early, and was very severe and long-lasting. Various accounts state that the ground was frozen 
with hard ice from 2 October until the end of April 1784. All fiords were reported to be frozen 
over in late February 1784 (for the first time in 39 years), and sea ice was very widespread and 
long-lasting. The summer of 1784 was also cold and wet, with occasional periods of frost and 
sleet. Ogilvie’s summary of seasonal weather shows that there was uniformly cold weather across 
Iceland for five seasons (Summer 1783 through summer 1784) after the onset of the Laki 
eruption. Also, 1782 was unusually cold, although apparently not as uniformly so as following 
the eruption. 


Haze and Dust 

The tremendous quantity of volcanic gases and dust released during the Laki eruption was 
reported from many locations in the northern hemisphere. The English rector Gilbert White 
(1977) wrote: 

The summer of the year 1783 was an amazing and portentous one, and full of 
horrible phenomena...the peculiar haze, or smoky fog, that prevailed for many 
weeks in this island, and in every part of Europe, and even beyond its limits, was 
a most extraordinary appearance, unlike anything known within the memory of 
man. ...The sun at noon looked as blank as a clouded moon, and shed a rust- 
coloured ferruginous light...and was particularly lurid and blood-coloured at rising 
and setting. 

Icelandic accounts describe the volcanic haze from early June to the end of August (Ogilvie 
1986), and Lamb (1970) reported the following first sightings of dry fog or haze in Europe and 

Copenhagen 29 May 
France 6 June 
North Italy 18 June 
Syria 1 July 
Altai, central Russia 1 July. 

Evidently, the dry fog spread eastward and southeastward at an average rate of approximately 
250 km/day during the first month of the eruption. This is only about 10% of the rate of 
propagation for the Krakatau haze (2,700 km/day; Russell and Archibald 1888). The difference 
in velocity may be due to the differences in direction (east for Laki, west for Krakatau) or the 
altitude (tropospheric for Laki, stratospheric for Krakatau). Benjamin Franklin (1784) also noted 
that the haze was seen in North America, although the original sources and details of this 
observation were not reported. Nonetheless, the haze persisted long enough, or rose to different 
atmospheric heights with differing wind directions, so that it was carried both eastward to Europe 
and westward to North America. 

Volcanic dust also fell out of the sky in Europe. Lamb (1971) reported that tulips in Holland 
were damaged by the dust and sulphurous smells during 18-24 June 1783. In Scotland the dust 
was thick enough to destroy crops in June. The detection of sulphurous odours in Europe proves 
that the haze was volcanic and not from some unknown forest fire, for example. The odour and 
eye irritation imply that the haze was at low altitudes. Volcanic dust that fell in Holland 11 days 
after the eruption started, apparently was transported much more slowly than dust from other 
Icelandic eruptions. Ash from the 1875 eruption of Askja reached Europe within a day 
(Thorarinsson 1963). 

The Summer of 1783 
White’s (1977) account quoted above continues: 
All the time the heat was so intense that butchers’ meat could hardly be eaten on 

the day it was killed; and the flies swarmed so in the lanes and amid hedges that 
they rendered the horses half frantic, and riding irksome. 


Instrumental temperature records reveal that 1783 was the warmest English July on record 
(Kington 1978). Other early thermometer data for six other major European cities allow 
quantification of how extreme the summer heat was in 1783. World Weather Records data 
(Figure 2) for Stockholm, Copenhagen, Edinburgh, Berlin, Geneva, and Vienna for a 31-year 
period centred on 1783 demonstrate that July 1783 was 1.6 to 3.3°C warmer than the 31-year 
average. In general the amount of the July temperature anomaly is closely correlated with the 
distance of each city from Laki (Figure 3). Thus, however the haze raised the summer 
temperatures in Europe, the effect was most pronounced where the haze was thickest, and the 
excess heating declined where the haze was probably less intense. Temperature data from eastern 
North America (Landsberg et al. 1968) reveal that the summer of 1783 was significantly hotter 
than the 225-year average (Sigurdsson 1982). Figure 4 shows the same data graphically. 





31 yr Average — 



Average Temperature (°C) 

1778 1780 1782 1784 1786 1788 


Figure 2: Average July temperatures (data from World Weather Records) for six European cities 
(Stockholm, Copenhagen, Edinburgh, Berlin, Geneva, and Vienna) for the seven-year period 
centred on 1783, the year of Laki’s eruption. The 31-year average is based on recorded 
temperatures for the 31 years centred on 1783. 




27 Edinburgh 


1000 2000 3000 

Laki to City Distance in Km 

July T Anomaly (degrees C) 

Figure 3: Deviation of July 1783 temperatures from 31-year averages as a function of the distance from 
Laki to six European cities. Edinburgh’s anomaly is less than expected based on the other 
cities, suggesting that the temperature increases were not latitudinally uniform. Perhaps these 
anomalies - from only one month after the start of the eruption - were due to tropospheric dust, 
which would not be as uniformly distributed as stratospheric aerosols. 

There are various sources of proxy weather information for this period; e.g., in Switzerland the 
summer was drier as well as warmer than normal (Pfister 1981), and there was a drought and 
poor harvest in Finland (Schove 1954). 

Additional circumstantial evidence that the summer of 1783 was warm includes a severe drought 
in the Yangtze region of China (Wang and Zhao 1981). The Yangtze drought continued into 
1784, but in both years there were floods in southeastern China and the Hwang Ho (Yellow) 
River Basin. An extraordinarily severe famine throughout Japan in the summer of 1783, however, 
was not caused by drought: Mikami (1988; Mikami and Tsukamura, this volume) has shown that 
an excess of rain destroyed many crops, and that, in fact, the summer of 1783 was wettest and 
coolest in Japanese history. Based on the high price of wheat in Delhi, India, the rains probably 
failed in 1783 with a consequent famine (Pant et al. 1988). These extremes in Asian weather 
during the summer of 1783 exhibit regional variations in the response to volcanism that are 
similar to previously-documented patterns in North America (Lough and Fritts 1987). 



Sigurdsson, 1982 

ee ae Variance -- 


Summer T (°C) 

— 225 yr Average 

1782 1783 1784 1785 1786 


Figure 4: Average summer temperatures in the eastern United States in the 1780s compared to the 225- 
year average. Data from Landsberg et al. (1968) as reported by Sigurdsson (1982). 

Winter of 1783-84 

Scant temperature measurements, abundant proxy data and anecdotal accounts demonstrate that . 
the winter of 1783-84 was one of the most severe on record in Europe and North America. The 
average January temperature for six European cities was 3° below the 31-year average centred 
on 1784 (Figure 5). Proxy temperature data (from viticulture/agriculture) indicate Switzerland 
had two extremely severe winters in 1783-84 and 1784-85, with the first year being the worst 
(Pfister 1981). The longest period of sea ice around Iceland also occurred during the winter of 
1783-84 when temperatures were nearly 5° colder than the 225-year average (Sigurdsson 1982). 
The next two winters were also significantly colder than normal (Figure 6). Information compiled 
by Ludlum (1966) includes the following records for the winter of 1783-84 in the eastern United 

Longest in early American history (last snow in late April), 

Near record depth of snowcover, 

Near record low temperatures, 

Greatest seasonal snowfall ever in New Jersey, 

Longest period of below zero temperatures ever in New England, 


Longest freezing ever of Chesapeake Bay, 

Longest and coldest winter in Maine, 

One of the greatest southern snowstorms (18-19 December), 

Freezing of Charleston Harbour (ice skating occurred), 

Freezing of Mississippi River at New Orleans (13-19 February 1784), 
Ice floes in Gulf of Mexico 100 km south of New Orleans. 

QO 2 
= awiG 
eh el 
® -2 
o 4 
= ee 
1778 1780 1782 1784 1786 1788 

Figure 5: Average January temperatures (data from World Weather Records) for six European cities 
(Stockholm, Copenhagen, Edinburgh, Berlin, Geneva, and Vienna) for the seven-year period 
centred on 1783, the year of Laki’s eruption. The 31-year average is based on recorded 
temperatures for the 31 years centred on 1783. 


Winter T (°C) 

1780 1781 1782 1783 1784 1785 1786 

Figure 6: Average winter temperatures in the eastern United States in the 1780s compared to the 225-year 
average. Data from Landsberg er al. (1968) as reported by Sigurdsson (1982). 

Ludlum (1966) provides graphic evidence of the severity of the winter by quoting contemporary 
letters and newspapers. Following a series of early and frequent storms, the worst weather of the 
winter occurred in mid-February, when minimum recorded temperatures for eight nights at 
Hartford, Connecticut were about 12°C or colder. From Virginia, James Madison wrote on 
11 February 1784, "We had a severer season and particularly a greater quantity of snow than is 
remembered to have distinguished any preceding winter." In another letter of 5 March 1784, 
George Washington complained that he, “arrived at this Cottage [Mount Vernon, Virginia] on 
Christmas eve, where I have been locked up ever since in frost and snow." 

The February cold spell froze the western end of Long Island Sound, and at New York City the 
Narrows between Staten Island and Long Island were blocked by ice for 10 days, preventing 
ships in Manhattan harbours from reaching the sea. Baltimore harbour was frozen by 2 January 
1784 and remained closed until 25 March. Chesapeake Bay was nearly completely frozen, and 
the Delaware River at Philadelphia froze on 26 December 1783 and was icebound until 12 March 
1784. Ludlum reports that even the southern harbour of Charleston, South Carolina was frozen 
in February, “having produced ice strong enough for skating on, which is very uncommon there." 
The most amazing phenomenon of the winter was the freezing of the Mississippi River at New 
Orleans, which Ludlum (1966, p. 154) reports has happened only once before (1899): 


On the 13th of February, 1784, the whole bed of the river, in front of New 
Orleans, was filled up with fragments of ice, the size of most of which was from 
twelve to thirty feet, with a thickness of two to three. This mass of ice was so 
compact, that it formed a field of four hundred yards in width, so that all 
communications was interrupted for five days between the two banks of the 
Mississippi. On the 19th, these lumps of ice were no longer to be seen. "The 
rapidity of the current being then at the rate of two thousand and four hundred 
yards an hour," says Villars, “and the drifting of the ice by New Orleans having 
taken five days, it follows that it must have occupied in length a space of about one 
hundred and twenty miles. These floating masses of ice were met by ships in the 
28th degree of latitude [in the Gulf of Mexico]. 

That the unusually cold winter was not just confined to the eastern United States is clear from 
the Hudson’s Bay Company records indicating 1783-84 had the fifth worse ice blockage of 
Hudson Strait on record (Catchpole 1988). 

Summer of 1784 

Summer temperature in England averaged 0.5°C, and as much as 1.6°C, below the long-term 
norm during 1784. The driest 12 months in English history began in August 1784 (Kington 
1978). Temperatures in the eastern United States tended to be below average: <53.9°F (12.1°C) 
in Philadelphia and <48.0°F (8.9°C) in New Haven (Bray 1978). Tree-rings indicate a marked 
growth minimum in the growing season of Douglas fir in Nevada, Utah, and Wyoming in 1784 
and 1785 (Woodhouse 1988). Tree-ring densities from the Mackenzie Delta region of Canada 
indicate that 1784 had a very cold summer (Parker 1988). Light coloured rings in black spruce 
at the treeline near Quebec indicate low temperatures shortened the growing season in 1784 
(Filion et al. 1986). Similarly, tree rings from Alaska show that the cool weather of 1784 
extended far to the north (Oswalt 1957). 

Winter of 1784-85 

In Switzerland, the long duration of snowcover during the 1784-85 winter resulted in widespread 
growth of the snow mold Fusarium nivale, which led to harvest failure of the spring grain crops 
(Pfister 1981). In Bern the winter was also very severe, with snow on the ground for more than 
150 days (Pfister 1978). Winter, spring, and early summer of 1784 in Brittany were disastrous, 
with a cold winter, hail at the end of April, spring floods, and a drought until the end of June 
(Sutherland 1981). In England, 1785 tied with 1674 as the coldest March on record (Kington 
1978). The date of freezing of Lake Suwa, Japan (Figure 7), occurred 22 days earlier than the 
long-term average (Arakawa 1954). 

Subsequent Seasons 

1785 was the worst year of the decade in Brittany, and one of the worst of the century. In some 
areas no rain fell between January and August (Sutherland 1981). The summer was also cool in 
England (Bray 1978), and the autumn of 1786 was one of the three coldest in English history 
(Kington 1978). 

In the eastern United States, summer temperatures were lower than normal in New Haven in 
1785 but returned to normal in 1786 (Bray 1978). 


Arakawa (1954) 

Feb. 1 

Jan. 1 

Lake Freeze Date 

Dec. 1 
1770 1780 1790 1800 


Freeze Date (Days Earlier 
or Later Than Average) 

Figure 7: 23-year record of the date of freezing of Lake Suwa, Japan centred on 1783 (data from 
Arakawa 1954). In 1783 freezing occurred 22 days earlier than normal. 

Mechanisms to Explain Observed Climatic Anomalies 

Each of the unusual weather records reported above can be dismissed as a freak occurrence 
within the normal range of variation, and thus not requiring a special origin. Consideration of 
the entire list of anomalies - and these are only the items that were found during a brief 
examination of secondary and tertiary historic records - suggests, however, that a period of 
unusual weather affected various places in the northern hemisphere from the summer of 1783 
through 1785. The principal observations to be accounted for are: 

Early summer 1783: Dry fog over Europe, western Asia, and the United States 
Summer 1783: Hot in Europe, United States and China; cold in Iceland 
Winter 1783-84: Exceptionally cold in Europe, United States and Japan 

Winter 1784-85: Very cold in Europe and Japan 

Summer 1785: Cool and dry in Europe and United States 


Benjamin Franklin’s famous 1784 communication to the Literary and Philosophical Association 
of Manchester was the first suggestion that volcanic eruptions might effect climate (see 
Sigurdsson 1982): 

During several of the summer months of the year 1783, when the effect of the 
sun’s rays to heat the earth in these northern regions should have been greatest, 
there existed a constant fog over all of Europe, and a great part of North America. 
This fog was of a permanent nature; it was dry, and the rays of the sun seemed to 
have little effect toward dissipating it, as they easily do a moist fog, arising from 
water. They were indeed rendered so faint in passing through it, that when 
collected in the focus of a burning glass, they would scarce kindle brown paper. 
Of course, their summer effect in heating the earth was exceedingly diminished. 

Hence the surface was early frozen. 
Hence the first snows remained on it unmelted, and received continual additions. 
Hence the air was more chilled, and the winds more severely cold. 

Hence perhaps the winter of 1783-84 was more severe, than any that had happened 
for many years. 

The cause of this universal fog is not yet ascertained. Whether it was adventitious 
to this earth, and merely a smoke, proceeding from the consumption by fire of 
some of those great burning balls or globes which we happen to meet with in our 
rapid course around the sun, and which are sometimes seen to kindle and be 
destroyed in passing our atmosphere, and whose smoke might be attracted and 
retained by our earth: or whether it was the vast quantity of smoke, long continuing 
to issue during the summer from Hecla in Iceland, and that other volcano which 
arose out of the sea near that island, which smoke might be spread by various 
winds, over the northern part of the world, is yet uncertain.’ 

Franklin’s first speculation, that the summer fog and winter coldness could be due to smoke from 
a meteor is rather bizarre, and has been forgotten (except perhaps as an unremembered 
contribution to the idea that a comet or asteroid collision with the Earth 65 million years ago 
resulted in extinction of the dinosaurs). His second option, that smoke from “Hecla in Iceland, 
and that other volcano which rose out of the sea near that island" caused the observed weather 
anomalies is more enduring. 

The general concept that volcanic activity can affect the climate has been developed since the 
obvious weather anomalies following the eruption of Krakatau in 1883 (Russell and Archibald 
1888). Mitchell (1961); Lamb (1971); Self et al. (1981); Devine et al. (1984) and others have 
demonstrated statistically that years in which volcanic aerosols are ejected into the stratosphere 
by volcanic eruptions are typically followed by one to three years of temperatures 0.2 - 0.5°C 
below average. Thus, although there are doubters, there is a widely-promoted model that 
explosive eruptions of sulphur-rich magmas can implant sulphuric-acid aerosols into the 
stratosphere where they remain suspended for a few years. The aerosols spread around the globe, 

' Ttalics added by editor. 


absorbing incoming solar radiation, thus heating the stratosphere and, by a reduction in the 
amount of radiation reaching the ground, cooling the Earth’s surface (Sigurdsson 1982; Rampino 
and Self 1984). 

The Laki eruption does not appear to fit this general model because it was not a classical 
eruption, such as Krakatau or Tambora, which explosively ejected aerosols into the stratosphere. 
Effusive eruptions, like Laki and typical of activity at Hawaiian volcanoes, are thought to have 
only minimal explosive activity, with nearly all magma flowing quietly across the Earth’s surface 
as lava flows. Wood (1984a); Devine et al. (1984); and Stothers et al. (1986) proposed, however, 
that aerosols from Laki may have entered the stratosphere even though the eruption was largely 
quiescent. The last two groups proposed that heat from fire fountains and lava fields could 
generate convective plumes that would rise into the stratosphere. This effect would be enhanced — 
by normal atmospheric mixing across the tropopause which replaces the entire air mass in the 
lower and middle stratosphere with tropospheric air every one to two years (Flohn 1968). 
Following my previous suggestion (Wood 1984a), this mixing could result, over the prolonged 
period of the Laki eruption and with the possibly high altitude of convectively-transported 
materials, in substantial deposition of sulphur aerosols in the stratosphere. Thus, the Laki 
eruption may have emplaced sufficient material in the stratosphere to produce the multi-year 
climatic effects. 

Most of the observed climatic effects in the early to mid-1780s can be explained by the volcanic 
hypothesis (Lamb 1970). The dry fog in Europe, the Near East, and North America, and the 
sulphurous smells, burning of eyes, and singeing of tulips in western Europe resulted from the 
tropospheric movement of Laki dust and sulphur aerosols mainly to the east but apparently also 
to the west. Gilbert White reported that the dry fog lasted one month, which coincided with the 
period of maximum volcanic activity (Wood 1984a). The hot summer weather in 1783 in Europe, 
United States and China is an unusual occurrence; no other volcanic eruption is associated with 
such hot weather. Perhaps the heated gases in the mid-troposphere hindered normal convection 
so that heat was trapped near the surface. The cold summer in Iceland, however, was presumably 
caused by the blockage of sunlight by persistent dense haze and smoke from Laki. Similar 
immediate cooling near dense volcanic plumes occurred at Tambora (Rampino, this volume), as 
well as Krakatau and Mount St. Helens (10° and 8°C below normal, respectively; reported in 
Simkin and Fiske 1983). 

The exceptionally cold winter of 1783-84 in Europe, North America and Japan is proposed to 
have resulted from the standard volcanic mechanism of stratospheric warming and hence 
tropospheric cooling due to the abundance of volcanic sulphuric-acid aerosols in the stratosphere. 
The very cold winter in Europe and Japan during 1784-85 and the following (1785) cool and dry 
summer in Europe and United States can only be explained if large numbers of aerosols remained 
in the stratosphere through 1785. This would be consistent with other large explosive eruptions, 
such as Tambora and Krakatau which were followed by lower than normal temperatures for one 
to three years (Rampino and Self 1984). 

These explanations would be impossible if aerosols from Laki did not enter the stratosphere. One 
significant piece of evidence suggests they did not. Sulphuric acid droplets from volcanic-eruption 

clouds fall to the Earth everywhere under the passing cloud, but the existence of the droplets is 
recorded only when they fall on permanent ice fields, as in Greenland and Antarctica, where they 


leave an acidic trace in that year’s ice layer (Hammer 1977). The largest acid spike in the Camp 
Century ice core in Greenland occurs in 1783 (Hammer 1977); but as pointed out by Sigurdsson 
(1982) there is no acid anomaly for 1784. The 1783 anomaly could be due to either tropospheric 
or stratospheric transport of aerosols from Laki (only about 1200 km to the east). An anomaly 
for 1784, which could only occur if significant aerosols were stored in the stratosphere for a 
year, would be strong evidence that Laki materials reached the stratosphere; the lack of a 1784 
anomaly is most consistent with no stratospheric contribution. If this is true then the present 
understanding of volcanic influences on climate require that the cold winter of 1784-85 is 
unrelated to the Laki eruption and the cold winter of 1783-84. 

A second uncertainty is whether Laki was actually the volcano that produced the anomaly 
recorded in the Camp Century ice core and that caused the observed climatic effects. One 
confusing piece of evidence is the report (in Lamb 1970) that the dry fog was first observed on 
29 May 1783 in Copenhagen and on 6 June in France. Yet the Laki eruption is recorded to have 
started only on 8 June 1783. Was another, earlier eruption responsible for the dry fog and other 

Eleven eruptions are recorded to have begun in 1783 (Simkin ef al. 1981), and two other little 
known ones may have occurred (Table 1). In terms of volume, Laki was the largest eruption and 
Asama, in Japan was the second largest. All other eruptions of the year are thought to have been 
considerably smaller, but one of them may have been important. 


As Franklin (1784) noted there was another "...volcano which rose out of the sea near..." 
Iceland. This volcano was the temporary island of Nyey or Noyde (New Island) which formed 
over the Mid-Atlantic Ridge some 50 km southwest of the Reykanes Peninsula in southwestern 
Iceland. Nyey began erupting by 1 May 1783 and produced a large deposit of pumice that floated 
on the sea for about 250 km around the volcano, causing great hardship for sailors (Lyell 1969). 
By autumn, when the Danish government sent an expedition to lay claim to the island, Nyey had 
been destroyed by wave action. One of the few descriptions (and a drawing, Figure 8) of the 
eruption is reproduced in Thorarinsson (1967). The Danish Captain Mindelberg of the brig 
Boesand first saw a smoke column on | May and wrote in his ship log, "At three o’clock in the 
morning we saw smoke rising from the sea and thought it to be land; but on closer consideration 
we concluded that this was a special wonder wrought by God and that a natural sea could 
burn...When I caught sight of this terrifying smoke I felt convinced that Doomsday had come." 
(quoted by Thorarinsson 1967). On 3 May Boesand approached the area of the smoke plume, but 
...when we had come within half a mile of the island we had to turn away for fear that the crew 
might faint owing to the enormous sulphur stench." 

Two, perhaps similar, eruptions near Iceland during the last 25 years provide comparisons. In 
1963 the island of Surtsey formed off the southern coast, and in 1973, a new cone and lava flow 
was constructed near Surtsey at Heimaey on the Vestmann Islands. Both eruptions were similar 
to the account of Nyey in that explosive eruptions produced scoria cones, but both Surtsey and 
Heimaey were armoured by lava flows and have been able to withstand wave erosion. Neither 
of the recent island eruptions produced as large a pumice field as reported for Nyey, and only 
minor amounts of ash fell in Europe. Based on the modern examples, it seems unlikely that Nyey 
could have caused the widespread effects commonly attributed to Laki, but the reported 
widespread pumice and lack of detailed information makes it impossible to reject completely the 
notion that Nyey contributed to the 1783 climatic phenomena. 


Wi cat 
gos ie 

uw tik Gages Peitmnci, 
Li “gu & WAS segs = 

ee Pees 
Ai fk oo 


Figure 8: Drawing and last page of text from Captain Mindelberg’s report on the Nyey eruption 
southwest of Iceland in May 1783. Reproduced from Thorarinsson (1967). 


Table 1: Volcanic Eruptions, 1783. 

05 May Off Ieland 

09 May Asama Japan 4 Biggest eruption 
in August 

till 8 Feb. 1784 

ZL Seo eee 
2 a! 

03 December 

' From Simkin et al. (1981). ? = Unknown date within 1783; ?? = uncertainty if eruption 
occurred in 1783. VEI 4 = Volcanic eruption index (0-8); VEI 4 = 10° to 10° m’ of ejecta. 


The largest historic eruption of Asama volcano in Japan began on 9 May 1783. Bullard (1976) 
and others have suggested that this eruption caused the climatic anomalies of 1783. I have 
previously summarized (Wood 1984b) recent Japanese literature on the 1783 Asama eruption 
which tends to discount it as the source of the dry fogs and other early summer climatic effects. 
The main argument is that although eruptive activity began in early May, nearly half of the total 
of 0.5 km* of ejecta was deposited during two days of intense eruptions on 3 and 4 August 1783, 
and most of the remainder formed during the next five days (Imai and Mikada 1982), two months 
after the dry fogs were reported. Asama may have contributed to the generally cool winter of 
1783, but it did not contribute to the strong atmospheric effects of early summer. 


The observation that dry fog was reported in Europe 10 days before the onset of activity of Laki 
is well dated by eyewitness accounts. Thordarsson and Self (1988) discovered by studying old 
Icelandic maps that the Grimsvotn basaltic caldera, about 50 km northeast of Laki along the 
fissure trend, erupted repeatedly throughout the Laki eruption. As proposed by Sigurdsson and 
Sparks (1978), activity along the Laki fissure system was probably intimately tied to activity at 
the Grimsvotn caldera. Thordarsson and Self (1988) suggest that there may have been an eruption 
at Grimsvotn in May, before the first Laki activity. Thus, the Laki/Grimsvotn system may have 
produced all the dry fogs of the summer of 1783. 



There are many loose ends in the story of Laki and its possible climatic effects. In this report a 
variety of readily available observations of unusual climatic phenomena occurring during the two 
years following the eruption is presented. The simplest assumption is that these anomalies are 
related to Laki, just as similar types and durations of climatic phenomena are clearly accepted 
as being associated with Tambora’s eruption in 1815. It is most likely that eruptions of the 
Laki/Grimsvotn system caused the dry fogs and hot summer of 1783 and the cold winter of 1783- 
84. In order to cause the cold winter of 1783-84 volcanic aerosols must have reached the 
stratosphere. And probably the cold winter of 1784-85 was due to the same stratospheric aerosols, 
which however, left no trace in the Greenland ice core for 1784. If Laki produced all of these 
effects, present volcano-climate models are inadequate to explain how. If Laki was not . 
responsible, then a major eruption 200 years ago is completely missing from our records. 

In compiling the historical data for the 1780s it became obvious that most reports are from the 
eastern United States and western Europe. A much greater effort is required to search the 
historical (and proxy) archives of Africa, Asia, South and Central America, and central and 
western United States to further define possible climatic effects of Laki and other eruptions. 


I thank Michael Helfert for sharing information concerning the unusual weather following the 
eruption Laki. 


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The Effects of Major Volcanic Eruptions on Canadian Surface 

Walter R. Skinner’ 


The superposed epoch method of analysis was used to detect changes in Canadian surface 
temperatures due to large volcanic dust veils in the atmosphere. This method accentuates weak 
signals that are present in a data series, as a temperature signal caused by a volcanic dust veil is 
expected to be of the same magnitude as the background noise level. Lamb’s Dust Veil Index ~ 
(DVI), a measure of the amount of volcanic material injected into the atmosphere, was used to 
select the volcanic-eruption dates beginning with the eruption of Krakatau in 1883. The DVI is 
directly related to the total loss of solar radiation reaching the Earth’s surface, and has the 
advantage of not being calculated from temperature information. Surface-temperature records for 
up to 20 Canadian stations were analyzed on national, regional (Arctic) and seasonal (summer 
and winter) bases for both equatorial and mid-latitude eruptions. A small sample test of 
significance was applied, and all suspected temperature signals proved to be significant at the 
0.01 level or better. The annual temperature depression following a mid-latitude eruption was 
about 0.4°C, and occurred during the eruption year and lasted no longer. A decline in annual 
surface temperature of about 1.0°C occurred in the first year after an equatorial eruption and 
persisted to a lesser degree, for another year or so. The difference appears to be directly related 
to the substantially greater mean DVI for the equatorial eruptions. The annual temperature drop 
in the Arctic was slightly greater than that for the country as a whole. Summer-temperature 
signals were stronger than those in winter, and in almost all cases were of a greater magnitude 
than the annual signals. There was a marked drop in winter temperatures of about 1.0°C 
following an equatorial eruption. 


The eruption of El Chichén in southern Mexico between 28 March and 4 April 1982 ejected huge 
concentrations of gases and particles into the upper atmosphere. By mid-November 1982, detailed 
solar radiation measurements in Fairbanks, Alaska began to display distinct differences from the 
previous five-year normal (Wendler 1984). Clear days during the 15 November 1982 to 31 May 
1983 period, when compared to clear-day data for the previous five years, showed a decrease in 
the direct beam of almost 25% and a decrease in global radiation of about 5%. Mass and 
Schneider (1977) previously determined that large volcanic dust veils in the atmosphere can 
reduce direct solar radiation by as much as 10%. This is simultaneously accompanied by an 
increased scattering effect that could substantially change the total amount of solar radiation 
reaching the Earth’s surface. 

Many theoretical investigations (Schneider and Mass 1975; Pollack et al. 1976) and empirical 

studies (Lamb 1970; Oliver 1976; Mass and Schneider 1977; Taylor et al. 1980) have been made 
in an attempt to determine the possible influence of large volcanic dust veils on surface weather 

' Canadian Climate Centre, Atmospheric Environment Service, 4905 Dufferin Street, Downsview, Ontario M3H 5T4, 


and climate. Most of these investigations were conducted on either a global or hemispheric scale. 
Taylor et al. (1980) also searched for volcanic signals on latitudinal, continental/marine and 
seasonal bases. A drop in annual average surface temperature of between 0.5 and 1.0°C in the 
first or second year following a large volcanic eruption was found in most of the empirical 

Canada, having an extensive area in mid- and high latitudes, should experience volcanic dust veil 
influences at varying times after a major eruption depending upon both the location and the time 
of year of the eruption. Oliver (1976) estimated a mid-latitude eruption to have a same year 
impact on northern hemisphere mean temperatures, while a similar impact by an equatorial 
eruption would be delayed for about a year. Lamb (1970) states that the transfer of upper-level 
dust veils from equatorial to mid-latitudes is accomplished mainly in autumn, and to a lesser 
extent in spring, with the great seasonal circulation changes. 

In this investigation, surface-temperature records for up to 20 Canadian stations were analyzed 
on national, regional (Arctic) and seasonal (summer and winter) bases for both equatorial and 
mid-latitude eruptions. 

Methodology and Data 

The Superposed Epoch Method 

The superposed epoch method of analysis, as outlined by Panofsky and Brier (1965) and 
employed by Mass and Schneider (1977) and Taylor et al. (1980), was used to detect changes 
in Canadian surface temperatures due to large volcanic dust veils in the atmosphere. This method 
accentuates weak signals that are present in a data series. A temperature signal caused by a 
volcanic dust veil is expected to be of the same magnitude as the background noise level, or the 
variability of the atmosphere (Taylor et al. 1980). 

Volcanic Eruption Dates 

Volcanic eruptions were selected on the basis of amounts of material ejected into the atmosphere, 
latitude of the eruption and the isolation in time of the eruption from any other major volcanic 
event. Lamb’s (1970) Dust Veil Index (DVI) was used as a basis for selecting most of the 
eruption dates. This is a measure of the amount of volcanic material injected into the atmosphere, 
and is directly related to the loss of solar radiation reaching the Earth’s surface. The eruption of 
Krakatau in 1883 was given a value of 1000, and all other eruptions were adjusted to it. The DVI 
has the advantage of not being calculated from temperature information (Mass and Schneider 

Five of the six volcanic eruptions selected have a total DVI greater than 150, and are classed as 
major volcanic events (Table 1). The 1956 eruption was chosen because of its mid-latitude 
location and isolation in time from any other major volcanic event. Three equatorial and three 
mid-latitude eruptions were selected in an attempt to isolate both the temporal dimensions and the 
magnitudes of the temperature signals in Canada following a major volcanic event. 

Selected volcanic events had to be separated by at least five years from any other major volcanic 
event. This was done to avoid the problem of cumulative dust veils that might obscure resulting 
signals. The separation of the 1907 event from the preceding event (1902) and the following event 
(1912) is exactly five years. Incorporating the 1907 event was called for as it was a major 
mid-latitude event for which ample records were available. 


Table 1: Date, Location and Dust Veil Index (DVI) of Selected Major Volcanic Eruptions. 

# Eruption Location Dates Key Dates DVI DVI 
1. Krakatau, Indonesia 6.0°S 105.0° E Aug 1883 Aug 1883 1000 1000 
2. Mont Pelée, Martinique 15.0°N 61.0° W May 1902 - - 100 - 
Soufriére, St. Vincent 13.5°N 61.0°W May 1902 -— - 300 - 
Santa Maria, Guatemala 14.5° N 92.0° W Oct 1902 - - 600 - 
Cumulative Data: May 1902 - 1000 
3. Shytubelya Sopka, 
Kamchatka 52.0° N 157.5° W Mar 1907 Mar 1907 500 500 
4. Katmai, Alaska $8.0° N 155.0° W Jun 1912 Jun 1912 150 =6150 
5. Bezymjannaja, 
Kamchatka 56.0° N 160.5° E Mar 1956 Mar 1956 10 10 
6. Gunung Agung, Bali 8.5°S 115.5° E Mar 1963 Mar 1963 800 800 
Equatorial Eruptions (# 1, 2 and 6) Average Total DVI = 933 
Mid-Latitude Eruptions (# 3, 4 and 5) Average Total DVI = 220 

Eruption year key months were also determined from Lamb (1970). The key month was the 
month during which the volcano entered its most explosive phase. In the case of two or more 
eruptions in the same year, such as 1902, the month of the first eruption was used. Table 1 
includes the key eruption date for each selected event. 

Composite Key Dates and Composited Temperatures 

The key volcanic eruption date was defined as the 12-month period beginning with the month 
during which the eruption occurred. The use of this period results in a cleaner volcanic signal 
than using the actual calendar year of the eruption (Taylor et al. 1980). This 12-month period 
was termed the "eruption year", or year "0". Sequences of four preceding years, or the four 
12-month periods prior to the eruption year and the four following years, or the four 12-month 
periods after the eruption year, were then determined. These sequences provided the bases for 
both individual and multiple composites. The five annual periods, the eruption year and the four 
following years, were analyzed because 4 volcanic dust veil produced by a single eruption exists 
for only a few years (Lamb 1970). 

Average temperature values, for selected Canadian stations, were calculated for each month of 
each of the 12-month periods associated with an eruption year. The resulting 12 monthly values 
were then summed and averaged to yield an annual value for that particular year. Graphs based 
on individual volcanic events were then plotted and studied in an attempt to define climatic 


Annual temperature values for each individual eruption were then associated with the 
corresponding values for all other individual eruptions. In addition, values for equatorial 
eruptions were isolated and inter-associated. The same was done for mid-latitude eruption values. 
These corresponding values were then summed and averaged to yield a "superposed epoch". 
Graphs, based upon these multiple volcanic events, were plotted and analyzed in a comparable 
manner to the analyses of the individual events. 


The database used consisted of mean monthly temperature values for up to 20 Canadian weather 
stations over common time periods. Stations were selected on the basis of length and 
completeness of record and upon location. Thirteen stations were available for the 1883 eruption 
date. There were no long-term records available for this date west of Winnipeg. Four stations 
were added for the 1902 eruption date to provide east to west coast spatial coverage. Another 
station was added for the 1907 and 1912 eruptions. The lack of long-term records for stations in 
northern Canada restricts the study of the first four eruptions to more southerly Canadian 
latitudes. Northern stations were added for the last two eruptions. This brought the total to 20 
stations for the 1963 event. Table 2 shows the stations used for the 1883 eruption. Table 3 shows 
the stations added for the 1902, 1907 and 1912 eruptions. Table 4 lists the stations used for the 
1956 and 1963 eruptions. In some cases, such as Quebec City and Winnipeg, weather-observation 
sites were moved during the 1940s from city to airport locations. However, none of the eruptions 
used in this study occurred during this period. In addition, some long-term temperature records, 
such as those from Toronto and Montreal, have been subjected to an artificial warming due to 
the influence of urban expansion. It was hoped that this would have only minor influence on the 
results and that the method of analysis would subdue the apparent noise in this small portion of 
the data. 

Missing monthly values were estimated for each station by calculating the 30-year mean for that 
particular month. In most cases only one of the 13 to 20 values was absent. The resulting 
estimate had little effect on the overall monthly composite. There were never more than two 
missing values in any monthly composite. 

Canadian Analysis 

Taylor et al. (1980) found it necessary to use data from a group of stations rather than just 
individual stations when searching for a temperature signal related to a volcanic eruption. This 
is due to the year-to-year and station-to-station variability when dealing with single station 
superpositions. Thus, the superposed epoch method outlined previously was applied to some or 
all of the 20-station temperature database selected for this study. 

Individual Eruption Composites 

Figures 1 to 6 show the individual eruption dust veil temperature composites for the selected 
Canadian stations. The 1883, 1902, 1956 and 1963 composites each display a marked dip in 
average annual temperature either in the eruption year or in the following two years. The 1907 
and 1912 composites show no such dip during these years. The low values for the "-4" and "-3" 
years for the 1907 composite might be the result of a large 1902 dust veil. However, there is no 
such dip in the early years of the 1912 composite that might similarly be attributed to a 1907 dust 

Graphs, based on the multiple volcanic events, were then plotted and examined in an attempt to 
define volcanic signals. 


Table 2: Weather Stations Used in Studying the Influence of Volcanic Dust Veils on Canadian Surface 

Temperatures for the 1883 Krakatau Eruption. 

Weather Station 

1. Winnipeg, Manitoba 

2. Port Arthur, Ontario 

3. Ottawa, Ontario 

4. Beatrice, Ontario 

5. Woodstock, Ontario 

6. Toronto, Ontario 

7. Québec City, Quebec 

8. Montréal, Québec 

9. Chatham, New Brunswick 
10. Fredericton, New Brunswick 
11. Halifax, Nova Scotia 
12. Sydney, Nova Scotia 

13. St. John’s, Newfoundland 


49° 53’ N, 
48° 26’ N, 
45° 24’ N, 
45° 08’ N, 
43° 07’ N, 
43° 40’ N, 
46° 48’ N, 
45" 307.N; 
47° 03’ N, 
45° 57’ N, 
44° 39° N, 
46° 09’ N, 

47° 34’ N, 

97° 07’ W 

89° 13” W 

75° 43’ W 

79° 23’ W 

80° 45’ W 

79° 24’ W 



65° 29° W 

66° 36’ W 

63° 36’ W 

60° 12” W 

52° 42’ W 















Years AES No. 
67 5023243 
65 6046588 
64 6105887 
104. 6110605 
112. 6149625 
142 6158350 
88 7016280 
111 7025280 
75 8100990 
82 8101700 
63 8202198 
72 8205698 
83 8403500 

Table 3: Weather Stations Added to Those Used for the 1883 Krakatau Eruption for the 1902, 1907 and 
1912 Eruptions. 

Weather Station! Location Period Years AES No. 
1. Victoria, British Columbia 48° 25’ N, 123° 22’ W 1898-1981 84 1018610 
2. Medicine Hat, Alberta 50°01" N, 110943” W. 1883-1981 99 3034480 
3. Banff, Alberta S121? NSS 3470W, 1887-1981 95 3050520 
4. Regina, Saskatchewan 50° 26’ N, 104° 40, W 1883-1981 99 4016560 
5. Ottawa (CDA), Ontario 45° 23’ N, 75° 43’ W 1889-1981 93 6105976 

' Ottawa (6105887) not used for 1902 eruption. 


Table 4: Weather Stations Used In Studying the Influence of Volcanic Dust Veils on Canadian Surface 
Temperatures for the 1956 Bezymjannaja and 1983 Gunung Agung Eruptions. 

Weather Station! 














. Victoria, British Columbia 
. Medicine Hat, Alberta 

. Banff, Alberta 

. Regina, Saskatchewan 

. Winnipeg, Manitoba 

. Churchill, Manitoba 

. Ottawa CDA, Ontario 

Beatrice, Ontario 

Woodstock, Ontario 

. Toronto, Ontario 

Quebec City A, Quebec 
Montreal, Quebec 

Chatham A, New Brunswick 
Fredericton CDA, N.B. 
Halifax, Nova Scotia 
Sydney A, Nova Scotia 

St. John’s, Newfoundland 
Cambridge Bay, N.W.T. 
Mould Bay, N.W.T. 

Kuujjuag, Quebec 


48° 25’ N, 
SIS aN, 
50° 26’ N, 
49°53’ N; 
58° 45’ N, 
45°°23? N, 
45° 08’ N, 
43° 07’ N, 
43° 40’ N, 
46° 48’ N, 
45° 30’ N, 
47° O1N, 
45° 55’ N, 
44° 39’ N, 
46° 10’ N, 
47~ 35 N; 
69° 07’ N, 
76° 14’ N, 

58° 06’ N, 

12352272 WV. 

110° 43’ W 

115° 34’ W 

104° 40° W 

97° 07’ W 

94° 04’ W 

15° 437 W 

719° 23” W 

80° 45’ W 

19> 24° Wi 

ALL A23 OW, 

13° 35° W 

65° 27’ W 

66° 37’ W 

63° 34’ W 

60° 03’ W 

52° 44° W 

105° 01’ W 

119° 20’ W 

68° 25’ W 

St. John’s (8403501) not used for the 1956 eruption. 





































AES No. 





















Y 4.9 
w 4.7 
es ee: 
a 4.1 
od ata SS. 
ba} Sh 
ee a, 
fj Shq 3! 
Figure 1: 
rae Br We at 
4 Soe fe 5 
To 5 
HY 5.3 
E Sahel 
H 417 
c . 
me 4.5 
WY 4.3 
2 4.1 
SV als 
i Si5 TC 

KRAKATAU (1883) 

SNC Sr Wd 


Dust veil temperature composite. 
Thirteen Canadian station events. 


=) <3} os sl ot) ih OO 

Figure 2: Dust veil temperature composite. 



Powwow Us 

Seventeen Canadian station events. 


See er te) I ee a 

Figure 3: Dust veil temperature composite. 

Eighteen Canadian station events. 



KATMAL <191 2) 

ol ete i ees Wee PS ker 

Figure 4: Dust veil temperature composite. 


bee fe ee ee eC DO 0 oS 

Eighteen Canadian station events. 


<Ai=3) 2S Oe ee eased 

Figure 5: Dust veil temperature composite. 


— eee A) AINA 

Nineteen Canadian station events. 


=A: <3) 72) =O) teases 

Figure 6: Dust veil temperature composite. 


Twenty Canadian station events. 

The apparent significance of these graphs must be viewed with caution. The first four eruptions 
were embedded in a hemispheric-warming trend, whereas the last two eruptions occurred during 
a hemispheric-cooling trend (Mass and Schneider 1977). The year-to-year variability, or noise, 
found by Taylor et al. (1980) is evident in a Canadian context. The compositing of several 
volcanic events should reduce this noise level and accentuate a volcanic dust veil signal. 

Multiple Eruption Composites 

Figure 7 shows the temperature composite for all stations and all eruptions. There is an obvious 
temperature dip during the eruption year and the "+1" year. The temperature dip during these 
two years is about 0.4°C below the level of years "-4" to "-1". Figure 8 shows the composite 
for the three equatorial eruptions. There is a well-marked dip in the "+1" year, about 1.1°C 
beiow the level of years "-4 to "-1". Figure 9 shows the composite for the three mid-latitude 
eruptions. Here the temperature dip is in the actual eruption year, about 0.5° C below the levels 
of years "-4" to "-1". 

Arctic Analysis 

The data noise level, or year-to-year variability, when based upon different groupings of stations, 
should vary randomly while the volcanic signal should remain fairly constant (Taylor et al. 
1980). A regional analysis was one step in determining the significance of the possible volcanic 
signals outlined previously. It also provided the basis for volcanic signal investigation into a part 
of Canada which can be extremely sensitive to small alterations in surface temperature. 

The solar-radiation deficit produced by volcanic dust veils must be greatest in Arctic areas where 
dust veils persist longer and the sun’s rays travel obliquely through the layers of dust (Lamb 
1970). Reduced surface temperatures result in an accumulation of both sea ice and land snow. 
The increased albedo would produce a radiation deficit long after the dust veil has disappeared 
(Lamb 1970). It would also affect the general atmospheric circulation, possibly having 
far-reaching spatial effects. 

The superposed epoch analysis method was applied to four Canadian Arctic stations for the 1956 
and 1963 eruptions. There were no Canadian Arctic station records for the earlier eruptions. The 
stations used were Churchill, Manitoba, Cambridge Bay, Northwest Territories, Mould Bay, 
Northwest Territories and Kuujjuaq, Quebec. 

Figure 10 shows the average annual temperature composite for the 1956 and 1963 eruptions. The 
temperature dip in the "0" and "+1" years is similar to that in Figure 7 for all Canadian stations. 
It is about 1°C below the levels of years "-4" to "-1". The surrounding noise level, however, is 
quite different than that in Figure 7. Years "+2" to "+4" hint at Arctic temperature stability 
following a volcanic eruption. 

Seasonal Analyses 

Summer and winter investigations were made in an attempt to determine the relative magnitudes 
of the dust-veil signals. The key summer season was defined as the first three-month period (June 
to August) to follow an eruption. The key winter season was defined as the first three-month 
period (December to February) to follow an eruption. Sequences of four preceding and four 
following seasons were determined in the same manner outlined previously. Seasonal averages 
were calculated for all Canadian stations and for each year associated with a volcanic eruption. 




aa As 

a. 402 

4. { 


& 3.5 

[x] =f. oJ 0-2) =f) 2 ei a4 



Figure 7: Dust veil temperature composite. 
All Canadian station events. 



atyock} ch tl tye gh a 
Figure 8: Dust veil temperature composite. 

Fifty Canadian station events for 
three equatorial eruptions. 


ON) O = & GOIN 

=4) <3 <2) - 1s Oe Gi ise na 

Figure 9: Dust veil temperature composite. 
Fifty-four Canadian station events 
for three mid-latitude eruptions. 


1956. 1963 ERUPTIONS 

Se -10.5 
oe 10.7 
5 -11.3 
He -11.5 
me 711-7 
fx] -11.9 
Ay -12.1 
x -12.3 
be =Ate= Sea 2) el Owed sand oa 8 


Figure 10: Dust veil temperature composite. 
Canadian Arctic station events. 


Figures 11 to 13 show the summer season temperature composites. All but one graph show a 
distinct drop of up to several tenths of a degree in either the eruption year or the following year. 
These composites display a close resemblance to those in Figures 7 to 9. The magnitude of each 
temperature drop, however, is at least equal to or greater than that of the corresponding annual 


Figures 14 to 16 show the winter season temperature composites. There is a higher degree of 
year-to-year variability than there was in the summer composites. This makes it more difficult 
to detect a possible volcanic signal. There is a distinct drop in temperature during the first winter 
following an equatorial eruption (Figure 15). However, there is no such drop in temperature 
following a mid-latitude eruption (Figure 16). 

Significance Tests 

The fact that the regional Arctic analysis identified much the same volcanic signals as those of 
the national study is a supportive indication of significance. A more rigorous small-sample test, 
however, is desirable. 

The Student t-test was applied to the multiple eruption composites to determine whether the 
sample mean, or the mean of the one or two years during which the volcanic signal is evident, 
is significantly different than the population mean, or the mean of the nine years from which it 
was taken. Mass and Schneider (1977) applied this test to volcanic dust veil composites for 
northern hemisphere stations. The basic formula used was: 

where, x = sample mean 
pf. = population mean 
o = population standard deviation 
and, N = average number of stations, and x = number of eruptions in the composite. 






BG. 2 

Be 16.5 

= 16.3 

El 1604 

© 15.9 

is “4 -3s-2ueed Ce Ges a. gap 

Figure 11: Dust veil summer season temperature 
composite. All Canadian station 



ONwvrEF WUT e 

S4iP=3) <2, a1 Peal ec 4 

Figure 12: Dust veil summer season temperature 
composite. All Canadian station 



=qsgiese. cl 8) 2 sed 
Figure 13: Dust veil summer season temperature 

composite. All Canadian station 



1 ! 
‘0 be) . . 
mS £0 @) a ty Coe GC) C 

Sy Ry ee AN rds 

Figure 14: Dust veil winter season temperature 
composite. All Canadian station 


! 1 
3 ° . . 
oy a tpt ey 09 

aN ee} ery one a eg oe 

Figure 15: Dust veil winter season temperature 
composite. All Canadian station 




—() ce) cr} Si CL le 
Figure 16: Dust veil winter season temperature 

composite. All Canadian station 


Thus, for a mid-latitude composite for all stations 


18x3 = 54 

The degrees of freedom (Gregory 1963) are 

d.f. = (n-1) + (n-1) 


where, n, = number of population years 

Table 5: 

n, = number of sample years 

d.f. (1 sample year) = 8 
d.f. (2 sample years) = 9 

Figure Composite X 





All Stations 
All Events 3.70 

All stations 
Equatorial Events 2.91 

All Stations 
Mid-Latitude Events 3.68 

Arctic (a) -11.97 
(b) -12.15 


All Stations 

All Events 16.09 


All Stations 
Equatorial Events 15.93 

All Stations 
Mid-Latitude Events 15.95 

All Stations 
Equatorial Events -10.63 

Years pu 

(0,1) 3:91 
(1) 3.71 
(0) 4.13 
(0,1) -11.34 
(0) : 
(COFD)) 16257. 
(1,2) 16.48 
(0) 16.64 
(1,2) 3-9276 
























Student t-Test Calculations for Composites Having an Apparent Volcanic Signal. 









The problem encountered earlier concerning the low number of stations and eruptions when 
dealing with the Arctic composites needs to be discussed. The fewer the stations and eruptions 
used, the greater the difference between the means must be, in order to attain a given level of 
significance. Table 5 shows the calculated Student-t values and the associated levels of 
significance for all composites where a volcanic signal was apparent. The test results are similar 
to those of Mass and Schneider (1977). In all cases, there is a difference between the sample 
population and the entire set at a significance level («) of at least 0.01. 


Climatic variation is complex and influenced by many factors. It is therefore difficult to clearly 
identify possible volcanic influences. As a result, caution must be exercised when interpreting 
apparent historical evidence and using it to predict future events. However, the results of this 
study do provide some evidence of the effects of volcanic dust veils on surface temperatures. This 
allows some tentative conclusions to be made. 

The magnitude of the annual temperature drop, for all Canadian stations, was at least 0.5°C 
greater after the equatorial eruptions analyzed than after the mid-latitude eruptions analyzed. The 
average total DVI for the selected equatorial eruptions was 933, while it was 220 for the 
mid-latitude events. The mid-latitude temperature depression was about 0.4°C, occurring during 
the eruption year and lasting no longer. The equatorial signal of about 1.0°C occurred in the first 
year after the eruption year and persisted to a lesser degree, for another year or so. 

The annual temperature drop in the Arctic was slightly greater than that for the country as a 
whole. It was approximately 1.0°C, and occurred in both the eruption year and the year 
following. The lower significance levels for the Arctic signals reflects the small number of 
stations and events used. Further investigation of this region, using more stations and events, 
might be appropriate. 

Temperature signals were stronger in the summer than in the winter. In addition, the summer 
drops in temperature were, in almost all cases, of a greater magnitude than the annual drops. 
There was a marked drop in winter temperature of about 1.0°C in the year following an 
equatorial eruption. 

This investigation did not take trends of temperature into account. No technique, other than the 
compositing of several volcanic events, was used to eliminate trends. The first four eruptions 
selected occurred during hemispheric-warming trends, whereas the latter two occurred during 
hemispheric-cooling trends. An accurate assessment of volcanic dust veil signals would eliminate 
these trends before applying the compositing technique. The temperature results found in this 
investigation are quantitatively similar to the empirical results found by Mass and Schneider 
(1977) and Taylor et al. (1980) and to the theoretical results of Pollack et al. (1976). 


This project was undertaken in the Applications and Impact Division of the Canadian Climate 
Centre, Environment Canada. Mr. M.O. Berry provided project supervision. 



Gregory, S. 1963. Statistical Methods and the Geographer. Longmans, Green and Co. Ltd., 
London. 240 pp. 

Lamb, H.H. 1970. Volcanic dust in the atmosphere; with a chronology and assessment of its 
meteorological significance. Philosophical Transactions of the Royal Society, London 

Mass, C. and S.H. Schneider. 1977. Statistical evidence on the influence of sunspots and volcanic 
dust on long-term temperature records. Journal of Atmospheric Science 34:1995-2004. 

Oliver, R.C. 1976. On the response of hemispheric mean temperature to stratospheric dust: an | 
empirical approach. Journal of Applied Meteorology 15:933-950. 

Panofsky, H.A. and G.W. Brier. 1965. Some Applications of Statistics to Meteorology. First 
Edition. Pennsylvania State University. pp. 159-161. 

Pollack, J.B., O.B. Toon, C. Sagan, A Summers, B. Baldwin and W. Van Camp. 1976. 
Volcanic eruptions and climatic change: a theoretical assessment. Journal of Geophysical 
Research 81:1071-1083. 

Schneider, S.H. and C. Mass. 1975. Volcanic dust, sunspots and temperature trends. Science 

Taylor, B.L., T. Gal-Chen and S.H. Schneider. 1980. Volcanic eruptions and long-term 
temperature records: an empirical search for cause and effect. Quarterly Journal of the 

Royal Meteorological Society 106:175-199. 

Wendler, G. 1984. Effects of the El Chichén volcanic cloud on solar radiation received at 
Fairbanks, Alaska. Bulletin of the American Meteorological Society 65:216-218. 


Northern Hemisphere 

North America 

Climate of 1816 and 1811-20 as Reconstructed from Western North 
American Tree-Ring Chronologies 

J.M. Lough’ 


Reconstructed temperature and sea-level pressure anomalies are presented for the year 1816 and 
the decade 1811-20. The reconstructions were developed from western North American semi-arid 
site tree-ring chronologies. The reconstructed climatic conditions for North America and the 
North Pacific were not very anomalous for either 1816 or 1811-20. More unusual conditions were 
reconstructed in years other than 1816 between 1811-20, and for decades other than 1811-20 in 
the first half of the nineteenth century. The factors responsible for the unusual climatic conditions 
of the "year without a summer" do not appear to have affected surface climate of western North 
America to the extent that these conditions are translated into the climatic reconstructions. 


The exceptionally large eruption of Tambora in April 1815 has frequently been speculated to have 
been the cause of the unusual climatic conditions experienced in 1816 - “the year without a 
summer". Anomalous weather was recorded in that year in eastern North America and Europe 
(Milham 1924; Rampino and Self 1982; Stommel and Stommel 1983; Stothers 1984; and 
elsewhere in this volume). The extent of climatic anomalies outside of the regions bordering the 
North Atlantic has not, as yet, been appraised satisfactorily. 

Although empirical studies have provided evidence of large-scale area-averaged surface 
temperature decreases following major volcanic eruptions (e.g., Oliver 1976; Taylor et al. 1980; 
Self et al. 1981; Kelly and Sear 1984; Sear et al. 1987) and the results of a variety of models 
have supported the role of volcanic eruptions as a source of thermal forcing (e.g., Hunt 1977; 
Robock 1981; Gilliland 1982; Gilliland and Schneider 1984), the importance of volcanic eruptions 
(such as Tambora in 1815) as a major source of climatic variability is still disputed (e.g., 
Deirmendjian 1973; Landsberg and Albert 1974; Parker 1985; Ellsaesser 1986). Difficulties in 
assessing the role of volcanic eruptions in climatic variability arise for a number of reasons. 
Theoretical (e.g., Baldwin et al. 1976; Pollack et al. 1976) and empirical studies (e.g., Rampino 
and Self 1982, 1984) indicate that the amount of sulphate aerosols produced by an eruption is of 
more importance than the amount of silicate ash in determining the subsequent climatic impact. 
Unfortunately, most historical chronologies of volcanic eruptions (e.g., Lamb 1970; Hirschboeck 
1979-80; Newhall and Self 1982) do not provide measures of sulphate aerosols, only of the 
explosive magnitude of the eruptions, which is often assessed by the amount of ash produced. 
Acidity profiles from ice cores (e.g., Hammer et al. 1980; Legrand and Delmas 1987) can 
provide records of eruptions that produced considerable amounts of sulphuric acid aerosols. The 
ice-core records tend, however, to be biased towards eruptions occurring at higher latitudes at 
the expense of those occurring at lower latitudes, and so such records tend to be incomplete. 

Australian Institute of Marine Science, PMB 3, Townsville M.C., Queensland 4810, Australia. 


Other problems result from the small number of possibly climatically important volcanic eruptions ~ 
that have occurred during the period for which extensive instrumental climatic records are 
available. The small sample size limits the statistical inferences that can be made regarding the 
impact of volcanic eruptions on climate. Consequently, most empirical studies have examined 
temperature series averaged over zonal or hemispheric space scales and little attention has been 
given to the possible regional variations of a climatic response. For periods before the mid- 
nineteenth century, instrumental records can provide information for geographically limited 
regions, usually those bordering the North Atlantic (e.g., Angell and Korshover 1985). For 
periods prior to the introduction of widespread instrumental climatic records we must rely on 
proxy climatic information from documentary, geological and biological sources. Lough and 
Fritts (1987), for example, identified a possible spatial response of North American temperatures 
to low-latitude volcanic eruptions. The response comprised warming in the western states and 
cooling in the central and eastern states. This study was based on temporally and spatially detailed 
reconstructions of North American temperatures derived from western North American tree-ring 
chronologies, and covered the period from 1602 to 1900 A.D. Some verification of the 
reconstructed climatic response was provided by independent sources of proxy climatic 
information both within and outside of the study area. This is important as each proxy climatic 
record is an imperfect record of past climate. Each series contains bias and error terms which 
may be unrelated to climate. In addition, different series may respond to different climatic 
variables, in different seasons and with different frequency responses. The most comprehensive 
description and understanding of past climatic variations (and their possible causes) will, 
therefore, only be obtained by the careful comparison and integration of independent sources of 
information (e.g., National Academy of Science 1975; National Science Foundation 1987). 

As a contribution to the improved description and understanding of the climate of 1816 and the 
decade 1811-20, I present reconstructions of seasonal climate for North America and the North 
Pacific developed from western North American tree-ring width chronologies. 


The reconstructions used in this study were developed by H.C. Fritts and co-workers at the 
Laboratory of Tree-Ring Research, Tucson, Arizona, following the methods outlined by Fritts 
et al. (1979) and described in detail by Fritts (in press). Only a general description of some of 
the characteristics of these reconstructions is given here. Fritts (1976), Hughes et al. (1982) and 
Stockton et al. (1985) describe the general principles and procedures applied in 

An array of 65 low-altitude, semi-arid site tree-ring chronologies (Figure 1; Fritts and Shatz 
1975) was used to estimate, by canonical regression, seasonal values of temperature at 77 
stations, precipitation at 96 stations in the United States and southwestern Canada, sea-level 
pressure at 96 stations in the United States and southwestern Canada and sea-level pressure at 96 
gridpoints between 100°E and 80°W, 20°N and 70°N. Because of the general west to east 
movement of weather systems across North America it was possible to attempt reconstruction of 
climate outside the area covered by the tree-ring predictor grid (see also Kutzbach and Guetter 
1980). The temperature and precipitation models were calibrated over the period 1901-63, and 
those for sea-level pressure from 1899-1963. The temperature and precipitation estimates were 
verified with data independent of that used for model calibration. The general form of the final 
sea-level pressure models was verified using a subsample replication technique (Gordon 1982). 


The final reconstructions, representing the average of the two or three best-calibrated and verified 
models, were for each variable, station or gridpoint and season for the years from 1602 to 1961. 
The seasons were December to February (DJF), March to June (MAMJ), July to August (JA) 
and September to November (SON). The annual series were the average of the four seasonal 
reconstructions and, therefore, were from December to November. In retrospect, the use of the 
four- and two-month seasons has proved a drawback in comparing these reconstructions to other 
sources of information. 

The annual calibration and verification statistics can provide some insight into the reliability of 
these reconstructions. More than 30% of the temperature variance was explained over North 
America with values exceeding 50% over much of the central United States (Figure 2). Most of 
the region also showed reliability through positive reduction of error (RE) statistics and the 
majority of verification tests passed. Positive values of RE indicate that, over the verification 
period, the estimates are an improvement over simply assuming mean climatic conditions (Gordon 
1982). Areas of poor temperature reliability occurred in the northeastern United States, Florida 
and parts of Nevada and Colorado. 

Generally, less variance was calibrated for precipitation than temperature (Figure 2), with more 
than 30% variance explained only in an area extending along the eastern edge of the tree-ring 
predictor grid. Verification of the precipitation estimates was also poor over much of the region. 
The precipitation reconstructions appeared to be of much lower reliability throughout much of 
the southern and eastern United States. The canonical regression transfer function (which is based 
on matching of the large-scale patterns of climate and tree-growth represented by the major 
principal components of the respective grids) does not appear to be well-suited to the 
reconstruction of precipitation. This variable is dominated by small-scale processes and variability 
that are not well captured by this regression technique. This is despite the fact that the tree-ring 
chronologies used are most directly sensitive to precipitation (Fritts 1974). 

The calibration and verification statistics for sea-level pressure (Figure 3) show that more than 
30% of the variance was calibrated over a large part of the grid. The reconstructions tended to 
be least reliable over northeastern Asia - the area farthest removed from the tree-ring predictor 

Other general features of these reconstructions were (Fritts, in press): (a) sea-level pressure 
tended to be biased towards lower frequency climatic variations at the expense of high frequency, 
year-to-year variations; (b) autumn climate was poorly reconstructed for all three variables; 
(c) precipitation was least well reconstructed, and temperature was probably the most reliably 
reconstructed; (d) all reconstructions deteriorated in reliability downstream from the tree-ring 
predictor grid over eastern North America (where Atlantic influences outweigh those of the 
Pacific) and, for sea-level pressure, over eastern Asia; (e) the large-scale regional patterns of 
climatic variation were calibrated at the expense of precision at individual stations or gridpoints; 
and (f) the reliability of the reconstructions was enhanced by averaging over space and filtering 
through time. 





A 1500 or earlier 
® 1501 - 1600 


120° Oke 

Figure 1: Locations of 65 semi-arid site tree-ring chronologies in western North America. 


A full description of these reconstructions and their development is provided by Fritts (in press). 
The reconstructions have been applied in a number of studies into the nature of climatic variations 
in North America and the North Pacific and also compared with independent sources of climatic 
information (e.g., Fritts and Lough 1985; Gordon et al. 1985; Lough and Fritts 1985, 1987; 
Lough et al. 1987). These studies, together with analyses of the reconstructions themselves (Fritts 
in press), have provided insights into the strengths and weaknesses of this particular set of 
climatic estimates. In the words of H.C. Fritts (in press): "The specific conclusions regarding the 
climate from 1602 to 1960 are presented as tentative hypotheses derived from one dendroclimatic 
analysis and test. They must be compared to data from other independent paleoclimatic sources 
that can reveal changes on seasonal and decadal time scales with accurate yearly dates". 


DA ~ sh ae ai se 

FR ee a 



os * 

AYN ‘AI \\ 

er oe “Oy 

Nae ay \ 
40 Q \ S = YY” 
> i 

. gO 

50 ¢ 



sa ey ve 
= ee 

q ‘ 

Figure 2: Calibration and verification statistics for annual temperature (top) and annual precipitation 
(bottom). Percent variance explained with areas of greater than 30% shaded (left-hand 
figures); number of verification tests passed out of a total of five, and areas with RE statistic 
greater than zero shaded (right-hand figures). 

eae he, 

\\ - GG 



100E 120 140 160 80 > 

= pee 
: A rhe i 

é, pS 
\\) , x. 
f ay Jeanna Oh 

NNW ~ 
50 6 60 50 50 

100E 120 140 40 100 80W 

ee 46 y Nee a 2 Z 
an ame 5 a 

AS \\ 
an \ 275 

3.5 35 35 35 


Figure 3: Calibration and verification statistics for annual sea-level pressure. Notation as for Figure 2. 



The seasonal and annual reconstructed values of temperature and sea-level pressure for 1816 and 
1811-20 were compared with the reconstructed mean climate of 1901-60. The temperature 
reconstructions were standardized by the 1901-60 standard deviation (s.d.). The reconstructions 
and original 65 tree-ring chronology series were also compared with the mean of the whole 
period, 1602 to 1960, to assess how unusual 1816 and 1811-20 were in the longer-term context. 


The seasonal and annual reconstructions of temperature and sea-level pressure are presented in 
Figure 4. In the winter of 1815-16, the Aleutian Low was reconstructed to be displaced 
southeastwards, with slightly higher pressure reconstructed over the Canadian Arctic. 
Temperatures were reconstructed to be warmer in the western states (associated with enhanced 
southerly air flow) and cooler over the central and eastern states. Temperature departures up to 
2 s.d. below recent-period means were reconstructed over the Great Lakes. 

The large sea-level pressure anomalies reconstructed in spring over eastern Asia were in an area 
of low reconstruction reliability and were not, therefore, considered to be significant. The main 
reconstructed feature was a slight deepening of the Aleutian Low. The reconstructed temperature 
field did not exhibit very large anomalies, though temperatures were still warmer in the west and 
cooler in the central states compared to the 1901-60 normals. 

Discounting the sea-level pressure anomalies over eastern Asia, the reconstructed sea-level 
pressure field for summer did not show marked departures from the twentieth century mean 
values. Slightly lower pressure was reconstructed over western Hudson Bay. Temperatures were 
reconstructed to be slightly above the average over a large part of the United States, with below 
average conditions in the far western states. Temperatures were reconstructed to be close to the 
1901-60 mean over the northeastern United States, the area of extensively documented climatic 
anomalies for the summer of 1816. 

In autumn, a positive pressure anomaly was reconstructed over the eastern North Pacific that was 
linked with the colder temperatures reconstructed in the Pacific Northwest. Elsewhere in the 
United States, temperatures were reconstructed to be warmer than the 1901-60 mean values by 
up to 2 s.d. in the northeastern and southern states. However, the reconstructions are least 
reliable in autumn. 

In the annual average, discounting sea-level pressure anomalies over Asia, the major 
reconstructed feature was an area of higher pressure to the west of Hudson Bay. Higher sea-level 
pressure extended out over the Pacific, and lower pressure was reconstructed to the south. Thus, 
1816 seems to have been characterized by a weakened zonal circulation over the North Pacific. 
Temperatures were reconstructed to be warmer in the western states and cooler in the most 
southerly states. Although temperatures were reconstructed to be up to 1 s.d. below the 1901-60 
mean near the Great Lakes and northeastern United States, the main contribution to this appears 
to come from the temperatures reconstructed for the winter 1815-16. 


Figure 5 shows the reconstructed seasonal and annual sea-level pressure and temperature values 
averaged for the decade of 1811-20. The reconstructions were expressed as departures from the 
reconstructed mean of 1901-60, and those for temperature were standardized. 


a) DJF b) MAMJ 

c) JA d) SON 

\ Ss D, 

e) Annual 

Figure 4: -Reconstructed sea-level pressure (mb) and temperatures (s.d. units) expressed as departures 
from the 1901-60 means the year 1816 for: (a) winter; (b) spring; (c) summer; (d) autumn; 
and (e) annual data. 



a) DJF 

e) Annual 

Figure 5: Reconstructed sea-level pressure (mb) and temperatures (s.d. units) expressed as departures 
from the 1901-60 means for the decade 1811-20 for: (a) winter; (b) spring; (c) summer; (d) 
autumn; and (e) annual data. 


In winter, the Aleutian Low was reconstructed to be deeper than the average, with positive sea- 
level pressure departures over the Canadian Arctic. Temperatures were reconstructed to be cooler 
than the average through the central United States. 

In spring a negative sea-level pressure anomaly was reconstructed over Alaska with near-average 
conditions reconstructed elsewhere. Temperatures were reconstructed to be warmer than the 
average throughout most of the United States. These departures were significantly different from 
the 1901-60 mean, at the 5% level, for 73% of the 77 temperature stations. 

The summer sea-level pressure anomalies were reconstructed to be of small magnitude, with the 
exception of northeastern Asia. Temperatures were reconstructed to be cooler than average in the 
northwestern states and generally warmer than average in the central and eastern regions. There ~ 
was no evidence in these reconstructions of negative temperature anomalies in the eastern United 

The autumn sea-level pressure field was characterized by a positive anomaly in the northeastern 
North Pacific. Temperatures were reconstructed to be cooler than the average in the northwestern 
and western regions and warmer in the southeastern and eastern regions. 

In the annual average, the sea-level pressure anomalies (outside of Asia) were estimated to be 
of small magnitude. Slightly below average pressure was found over the North Pacific. 
Temperatures were reconstructed to be slightly warmer than the average over most of the United 
States, though at only 5% of the 77 stations were these values significantly different from the 
1901-60 mean values. 

Thus, the climate of 1811-20, as reconstructed from western North American tree-ring 
chronologies, did not appear to be particularly anomalous when compared to the mean climate 
of 1901-60. In the annual average, sea-level pressure was slightly lower and temperature slightly 
higher than the 1901-60 mean, but none of these departures was very large. 

The reconstructed climate of 1811-20 was compared with that reconstructed for the other four 
decades of the first half of the nineteenth century (Figure 6). These data were expressed as 
departures from the instrumental record mean of 1901-70, and precipitation was included, 
expressed as a percentage of the mean. In this context, 1811-20, appeared to have been the least 
unusual of the five decades. Extensive cooling was, for example, reconstructed in 1821-30, 1831- 
40 and 1841-50. Similarly, sea-level pressure anomalies of greater than 1 mb were evident in all 
decades except 1811-20. The climate as reconstructed from the western North American tree-ring 
chronologies for the decade 1811-20 was not very different from the recent mean conditions. 
More extreme climatic conditions were reconstructed for other decades in the first half of the 
nineteenth century. 

Comparisons with 1602-1960 Mean Conditions 

In the preceding sections the reconstructed climate of 1816 and 1811-20 was compared to recent, 
twentieth century mean conditions. The reconstructed climate did not appear to be very different 
from this mean. I examined the data with respect to the long-term 1602-1960 reconstruction 
mean. I also considered the nature of the anomalies of the ones tree-ring chronologies which 
were used to develop these climatic reconstructions. 



100E 120 140 16 180 160 140 120 100 BOW 


Figure 6: Reconstructed annual sea-level pressure (mb), temperatures (°C) and precipitation (percent of 
mean) expressed as departures from the 1901-70 instrumental record means for the first five 
decades of the nineteenth century. Dashed contour lines for the precipitation maps are through 
areas where the verification statistics indicate that the reconstructions are unreliable. 


The percentage of the 77 annual temperature stations and 65 tree-ring chronologies with 
departures of +1 s.d. and -1 s.d. of the 1602-1960 mean were calculated for each year of the 
decade 1811-20 (Table 1). The reconstructed temperature field was close to average conditions 
with only 3% of the stations with reconstructed values +1 s.d. of the mean in 1816. The years 
1811, 1818 and 1819 all had more than 45% of stations with departures +1 s.d. from the mean. 
In 1811, the departures were about equally above and below the mean, but in 1818 and 1819, 
they were mainly positive, indicating warmer conditions. 

Forty percent of the original tree-ring chronologies had departures of at least +1 s.d. of the 
1602-1960 mean in 1816, though this was not the most extreme year of the decade. The most 
extreme years were 1819 with 48% and 1818 with 42% of the 65 sites with departures +1 s.d. 
of the mean. For the last two years, the departures were mainly negative, indicating that 
conditions were generally unfavourable for tree growth. In contrast, in 1816, 35% of the 65 
stations had departures of at least +1 s.d. of the mean, indicating conditions were generally 
favourable for wide growth-ring formation in western North America. This was the most 
favourable year for the tree growth of the decade 1811-20. The term favourable for tree growth 
cannot be simply interpreted, as the 65 chronologies cover a range of tree species from different 
sites in western North America. Factors influencing the width of the annual growth ring vary 
considerably, and can also operate over a number of growing seasons (Fritts 1976). For semi-arid 
sites, wider annual rings are often, however, associated with moister and cooler conditions near 
the trees. 

The decade mean for each reconstructed variable and the tree-ring chronologies were compared 
to the long-term mean for 1602-1960 for each decade between 1602-10 (1602 was the first year 
of the reconstructions) and 1951-60 (Table 2). Evidently 1811-20 was not particularly unusual 
in these data. For temperature, 12% of the stations had departures significantly different from 
the long-term mean in 1811-20 compared to 62% of stations in the most extreme decade of 1681- 
90. For sea-level pressure, 1811-20 had 33% of the 96 gridpoints with significant departures 
compared to the most unusual decade of 1881-90 with 66%. None of the 96 precipitation stations 
was reconstructed to have values significantly different from the long-term mean in 1811-20, 
compared to 70% in 1611-20. For the original tree-ring chronologies, 19% were significantly 
different in 1811-20, compared to 60% for the most extreme decade of 1911-20. 


Table 1: Percentage of 77 Temperature Stations and 65 Tree-Ring Chronologies with Values + 1 Standard 
Deviation of 1602 to 1960 Mean for Each Year of the Decade 1811-20. 

Annual Temperature 

s.d s.d. d 
Year am! -1 | 
1811 29 25 eis) 
1812 27 10 37 
1813 2) 1 10 
1814 0 8 8 
1815 14 92 26 
1816 5) 0 3 
1817 10 0 10 
1818 47 0 47 
1819 38 8 46 
1820 26 a 33 

Tree-Ring Chronologies 

s.d. s.d. s.d. 

al -1 ae 
1811 15 6 a2, 
1812 9 6 15 
1813 6 23 29 
1814 12 11 23 
1815 14 9 23 
1816 35 5 40 
1817 22 5 27 
1818 11 Si 42 
1819 14 34 48 
1820 8 31 39 


Table 2: Percentage of Stations, Gridpoints or Chronologies for Which Decade Mean is Significantly | 
Different from Long-Term (1602-1960) Mean at the 5% Significance Level for Reconstructed 
Temperature (T), Sea-Level Pressure (SLP), Precipitation (PPT) and Tree-Ring Chronologies 


1602-1610 16 23 19 25 
1611-1620 40 21 70 40 
1621-1630 52 53 40 28 
1631-1640 22 50 52 20 
1641-1650 17 31 48 23 
1651-1660 12 52 5 19 
1661-1670 49 2 19 28 
1671-1680 39 46 5 22 
1681-1690 62 51 17 9 
1691-1700 1 11 4 14 
1701-1710 3 9 2 14 
1711-1720 0 15 0 9 
1721-1730 0 0 A 11 
1731-1740 0 16 3 28 
1741-1750 5 16 0 23 
1751-1760 9 1 30 20 
1761-1770 23 8 1 22 
1771-1780 30 71 1 26 
1781-1790 6 21 0 20 
1791-1800 31 17 7 23 
1801-1810 31 24 0 22 
1811-1820 12 33 0 19 
1821-1830 12 11 6 23 
1831-1840 39 18 41 42 
1841-1850 12 14 36 29 
1851-1860 0 8 1 11 
1861-1870 38 17 40 20 
1871-1880 3 ils} 7 25 
1881-1890 32 66 4 ie 
1891-1900 1 30 0 15 
1901-1910 6 52 4 17 
1911-1920 47 47 24 60 
1921-1930 19 5 5 35 
1931-1940 55 38 27, SZ 
1941-1950 14 36 21 31 
1951-1960 1S 48 - 39 

pe eS ee 


Summary and Conclusions 

As reconstructed by western North American semi-arid site tree-ring chronologies, the climate 
of North America and the North Pacific does not appear to have been very unusual in 1816 or 
the decade 1811-20. This is when compared to both the 1901-60 and the 1602-1960 reconstructed 
data means. 

For winter, spring and the annual average of 1816, temperatures were reconstructed to be cooler 
in the eastern and central United States and warmer in the western United States. In summer and 
autumn of 1816, temperatures were reconstructed to be warmer in the central and eastern regions 
and cooler in the west. The pattern of temperature departures for winter, spring and the annual 
average are similar to the average pattern identified by Lough and Fritts (1987) to characterize 
the years 0 to 2 after eight low-latitude volcanic eruptions between 1602 and 1900. The Tambora 
eruption of 1815 was one of eight eruptions used in that analysis. The summer temperature field 
for 1816 does not resemble the average pattern identified by Lough and Fritts (1987). In the 
annual average there was reconstructed to be a weakening of the westerly zonal flow pattern over 
the North Pacific. Sea-level pressure anomalies were not, however large. Thirty-five percent of 
the original tree-ring chronologies had growth departures of +1 s.d. or more above the 1602- 
1960 mean in 1816, indicating that conditions, at least in parts of western North America were 
generally favourable for wide tree-ring formation. 

The decade 1811-20, in the annual average, was reconstructed to be slightly warmer than the 
1901-60 mean over North America, with lower sea-level pressure reconstructed over the North 
Pacific. It was, perhaps, the least unusual of the first five decades of the nineteenth century. 
Relatively large negative temperature departures were reconstructed over the central northern 
United States in 1821-30, 1831-40 and 1841-50. The decade 1811-20 did not appear to be very 
unusual when compared to long-term mean conditions for any of the reconstructed variables nor 
the original tree-ring chronologies. 

The evidence from this particular set of climatic reconstructions from western North American 
semi-arid site tree-ring chronologies is for near-normal climatic conditions in 1816 and 1811-20. 
Reconstructed climatic anomalies were small in magnitude when compared to the recent, 1901-60, 
and long-term 1602-1960, mean conditions. Most references to the "year without a summer" in 
North America tend to come from eastern regions. Because this particular set of reconstructions 
is known to be less reliable in the east, where Atlantic and Arctic influences outweigh those of 
the Pacific, the lack of large reconstructed anomalies in this region was not surprising. What was 
Surprising was a lack of evidence for large-magnitude climatic anomalies in areas where the 
reconstructions are known to be reliable, over the western United States and the North Pacific. 
Analysis of the original tree-ring chronology series suggested that 1816 was a year favourable 
for tree-growth in parts of the western states, possibly associated with moister and cooler 
conditions. Large-scale climatic anomalies are not, however, apparent in the climatic 
reconstructions from these tree-ring data. This suggests that whatever the nature of the anomalies 
of climate in 1816 and the decade 1811-20, they were not large enough to significantly influence 
climatic conditions in the western United States either for good or bad. 


This study is based on the results of many years of work by Hal Fritts and co-workers at the 
Laboratory of Tree-Ring Research, University of Arizona. 



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Volcanic Effects on Colorado Plateau Douglas-Fir Tree Rings 

Malcolm K. Cleaveland' 

The explosion of Tambora in April 1815, is the largest volcanic eruption in recorded history. 
Based on measured temperatures, the significant North American climatic effects of Tambora 
appear to have been limited to the northeastern United States and eastern Canada. However, 
diameter growth of conifers on the Colorado Plateau in the southwestern United States was 
extremely large from 1815 to 1817, and the largest regionally-averaged late season (latewood) 
growth in 491 years occurred in 1816. This abnormal growth is probably not coincidental. 
Above-normal growth occurs when moisture stress is reduced, and the eruption probably resulted 
in abnormally low growing-season temperatures and/or abundant precipitation over the Colorado 
Plateau. Latewood density was above average from 1815 to 1817, also indicating reduced 
moisture stress. Trees located on marginal sites that are usually subject to the greatest moisture 
stress showed the most favourable growth response from 1815 to 1817. These growth changes 
are postulated to be effects of abnormally cool growing seasons that reduced evapotranspiration 
and delayed onset of drought-induced late summer dormancy. Delayed dormancy would favour 
development of an anomalously large latewood zone and increased latewood density. No 
comparable growth responses are apparent for other known large eruptions, indicating that 
regional climatic response to volcanic forcing is highly variable. Long-lived, climatically- 
responsive trees are widely distributed in the northern hemisphere. Analyses of these tree-ring 
data during recent centuries when instrumental climatic data are sparse may help reveal the 
impact of known volcanic eruptions on northern hemisphere climate, and may also help identify 
and date extremely large prehistoric eruptions. 


The April 1815 explosion of Mount Tambora at Latitude 8°S was the “largest and deadliest 
volcanic eruption in recorded history..." (Stothers 1984). Lamb’s (1970, Tables 7a,b) Dust Veil 
Index is larger for 1815 than any year since 1500. The 1815 eruption also has the highest 
Volcanic Explosivity Index since 1500 (Newhall and Self 1982). If such volcanic eruptions do 
affect global climate, then this huge eruption should have left evidence in instrumental climatic 
records and perhaps in other proxy climatic records such as ice cores and tree rings (Shutts and 
Green 1978; Bryson and Goodman 1980; Gilliland and Schneider 1984; Kelly and Sear 1984; 
Angell and Korshover 1985; Bradley 1988). 

Volcanic aerosols reduce solar radiation at the surface on a regional or global basis by increasing 
the albedo of the upper atmosphere. Temperature effects of aerosols are more closely related to 
the quantity of sulphates injected into the upper atmosphere than to the quantity of fine particulate 
ejecta. Sulphuric acid created as a result of sulphur-rich volcanic eruptions plays a major role in 
reducing transmission of direct solar radiation (Harshvardhan and Cess 1976; Pollack et al. 
1976). If the Tambora eruption had a global climatic impact, it must have created a sulphuric acid 
aerosol. In fact, the 1815 Tambora episode coincides with a very large acidity peak from 1815 

' Department of Geography, University of Arkansas, Fayetteville, Arkansas 72701, U.S.A. 

it Bs) 

to 1818 in Greenland ice cores, even before correction for losses in transport from equatorial to 
polar latitudes (Hammer et al. 1980; Rampino and Self 1984; Stothers 1984). 

Despite abundant evidence pointing to the 1815 Tambora eruption as having all the requisite 
characteristics for a major influence on global climate, Stothers (1984) and Angell and Korshover 
(1985) found little evidence of significantly lower surface temperatures following this eruption. 
Most of the available long instrumental-temperature records that form the basis of this conclusion, 
however, are confined to Europe and the northeastern United States. In addition, available 
records show that a period of well-below-average temperature started at least five years before 
1815, which may mask some of the volcanic effects (Stothers 1984; Angell and Korshover 1985; 
Baron, this volume). While 1816 is cooler than 1815 in most of these records, the drop is small 
compared to the general cooling shown in the period. The long temperature record from New 
Haven, Connecticut does show signs of considerable cooling, in keeping with the New England 
reputation of 1816 as "the year without a summer" (Stommel and Stommel 1983; Angell and 
Korshover 1985), but the northeastern United States and eastern Canada (Wilson 1985) are 
considered exceptions. Horstmeyer (1989) compiled Cincinnati, Ohio daily weather records from 
1814 to the present and found that, "No year since [1816] has even come close to having such 
a cold summer". This demonstrates for the first time that an abnormally cold summer of 1816 
occurred in the American Midwest, as well as in eastern North America. 

There is a large and growing network of old, climate-sensitive tree-ring chronologies available 
that might provide more spatially complete evidence for volcanic effects on climate (Stockton et 
al. 1985). One such network from western North America has been used to estimate seasonal and 
annual temperature variation for the United States (Fritts and Lough 1985; Lough, this volume). 
The reconstructed temperature estimates were then used to study average climatic response to 
selected volcanic events (1602 to 1900) by comparing temperature before and after the events 
with superimposed epoch analysis (Lough and Fritts 1987). After low-latitude eruptions like 
Tambora, cooling was especially pronounced in the spring and summer during the trees’ growing 
season. The spatial effects on United States climate were often different, and directly out of 
phase, between the west coast and the rest of the country. During the spring following eruptions, 
the Pacific Northwest experienced warming while the rest of the country cooled down. In summer 
the area of warming expanded down the west coast, while the central and eastern United States 
remained relatively cool. Lough and Fritts (1987), however, did not specifically report on the 
Tambora eruption. In this paper a set of climate-sensitive tree-ring width and density chronologies 
from the Colorado Plateau are used to investigate the possible impact of the 1815 eruption on the 
climate of the southwestern United States. 


Three sets of Douglas-fir (Pseudotsuga menziesii) tree-ring radial samples were collected and 
crossdated with standard methods (Stokes and Smiley 1968). The total number of radii in the 
collections ranged from 14 to 20, taken from 10 to 14 trees per site. The samples were collected 
at Ditch Canyon (DIT) on the Colorado/New Mexico border, at Mesa Verde National Park in 
Spruce Canyon (SPC) during 1978 (Cleaveland 1983, 1986, 1988) and in Bobcat Canyon (BOB) 
in 1972 (Drew 1976) (Figure 1). 



New Mexico 

Figure 1: Map of the three sites sampled: Bobcat Canyon (BOB), Spruce Canyon (SPC), and Ditch 
Canyon (DIT). 

Conifer tree rings are divided into earlywood and latewood zones, formed first and last in the 
growing season, respectively. Earlywood formation is most strongly influenced by climate in the 
spring and latewood formation by late spring to mid-summer climatic conditions. Typical 
earlywood cells become large, with relatively thin cell walls surrounding large cavities or lumens. 
Typical latewood cells are smaller than earlywood cells, with thick cell walls and small lumens. 
For this reason the earlywood part of a ring is less dense than the latewood portion. The 
transition to latewood is abrupt in Douglas-fir (Panshin and de Zeeuw 1970). The width of the 
two zones was measured optically from the BOB and DIT specimens. Characteristics of the SPC 
samples, including latewood width and average latewood density, were measured by X-ray 
densitometry (Parker et al. 1980; Cleaveland 1983, 1986). 

Time series of ring widths are often not statistically stationary because the mean and variance 
may both change with increasing age and diameter of the tree. The most common form of the 
growth function approximates an exponential curve declining to a constant value, but linear 
regression lines or more flexible polynomial or spline curves are also often used to remove 


growth trend (Stokes and Smiley 1968; Fritts 1976; Cook and Peters 1981). To transform the 
measurements into stationary time series, a curve is fitted to the measurements from each sample, 
and each annual value is divided by the corresponding annual curve value. This transforms 
measurement series into indices with a mean of 1.0, removing the effects of differences in mean 
growth from tree to tree, and rendering the variance quasi-stationary. The indices for each radial 
series from a site are averaged on an annual basis into a site chronology. The site chronology has 
a mean equal to 1.0 and a minimum value greater than 0.0, and represents a selected statistical 
sample of the macro-environmental factors that control the radial growth of a given species on 
a certain site through time. 

Results and Discussion 

BOB and DIT are lower forest-border sites that often experience high levels of moisture stress, 
whereas the SPC site is more mesic (Drew 1976; Cleaveland 1983, 1986). One measure of 
response to climate is the mean Sensitivity statistic, that is, the average first difference of 
chronology indices (Fritts 1976). The mean sensitivities of ring-width chronologies at BOB, DIT, 
and SPC are 0.45, 0.44, and 0.28, respectively. This statistic indicates that the BOB and DIT 
chronologies should show greater response to departures from normal growing-season conditions 
than the SPC chronology. 

A width or density index greater than 1.0 indicates above average growth that is usually 
attributable to a cool and/or moist growing season in the Southwest (Fritts et al. 1965; Fritts 
1976). When latewood width at the BOB, DIT, and SPC sites are averaged for each year, the 
average index is larger for 1816 than for any year since 1487, a period of almost five centuries 
(Figure 2). In addition, the ring-width, latewood-width, and latewood-density indices for the three 
collection sites all equal, or greatly exceed, average growth (1.00) for 1815, 1816, and 1817 
(Table 1). These anomalies indicate that the growing seasons were substantially cooler and/or 
wetter than normal (Cleaveland 1983, 1986). 

The very large values of latewood growth in the decade of the 1490s (Figure 2) are probably 
artifacts of a small number of samples, and end effects of curve fitting. The best replicated 
chronology, SPC, has an index of only 1.33 in 1491 (Figure 2). If the poorly-replicated 1490s 
are ignored, 1815 and 1816 summed are the largest average latewood total of two consecutive 
years, and 1816 and 1817 are the second largest. In addition, if the 1490s are not considered, the 
1815-17 period has the largest total latewood growth of three consecutive years. 

All chronologies are well replicated after about 1700, giving greater confidence in the estimated 
index means after 1700. It would certainly be possible to increase the sample depth of long series 
at many sites in the western United States to improve estimates in the early part of the 
chronologies. This should be an important consideration before using these chronologies to 
investigate earlier volcanic eruptions. 

Lower forest-border Douglas-fir trees in southwestern Colorado generally become dormant in 
June or July - forced into dormancy by moisture stress long before photoperiod or low 
temperatures could become responsible (Fritts et al. 1965). Also, a conditional probability 
analysis of 89 southwestern conifer chronologies (e.g., Stockton and Fritts 1971) indicates that 
the influence of temperature on tree growth late in the growing season is stronger than the 
influence of precipitation (Cleaveland, unpublished data). Chronologies at those sites showing the 
highest degree of inferred moisture stress (BOB and DIT) show a stronger response to the 


1815-17 climatic anomaly than the more mesic SPC site. It seems probable, therefore, that the 
growth anomaly is linked in some way to below-normal temperature and/or above normal 
precipitation that drastically reduced moisture stress on Colorado Plateau trees during those 
growing seasons. The greatly enlarged latewood zone in the 1815-17 rings of these Colorado 
Plateau conifers could be interpreted as evidence for a longer-than-normal growing season 
extended by below-normal air temperatures during the summer. Normal or above-average 
precipitation probably also occurred during the extended growing seasons from 1815 to 1817. 

Table 1: Southwestern Colorado Tree-Ring Chronology Indices (1810-20). 

Bobcat Canyon Ditch Canyon Spruce Canyon Average 
Ring Latewood Latewood Latewood Latewood* Latewood 
Year Width Width Width Width Density Width 
1810 0.68 0.70 Or 1.03 1.01 0.83 
1811 1.01 1.07 2.14 1.32 1.05 jlasyi | 
1812 fo12 0.87 1.09 0.87 1.01 0.94 
1813 0.41 0.36 0.75 0.31 0.88 0.47 
1814 0.92 0.92 1.03 0.69 0.97 0.88 
1815? 1.24 73 2.29 1.50 1.06 1.84 
1816 22 3.47 4.58 US 1.10 3.20 
1817 2.28 DoS 1.78 1.00 1.02 1.67 
1818 0.47 0.42 0.28 0.83 0.97 0.51 
1819 0.31 0.77 0.65 0.68 0.90 0.70 
1820 0.24 0.32 0.64 0.34 0.86 0.43 

Indices greater than 1.0 indicate above-average growth, and indices less than 1.0 represent 
below-average growth. 

Latewood density variability was multiplied by 3.0 to increase the range of variation relative 
to the other variables. 


The year Tambora erupted.Table 1 

Other historic eruptions are believed to have affected climate, and might have influenced the 
growth of trees on the Colorado Plateau. Rampino and Self (1984) list selected eruptions with 
estimates of the sulphuric acid aerosol generated and the estimated northern hemispheric 
temperature change. The eruption of Laki in 1783 is estimated to have caused greater cooling 
than Tambora, but no effect can be detected in Figure 2. Laki is a high-latitude (64°N) volcano, 
however, and Lough and Fritts (1987) found that volcanic eruptions in low latitudes resulted in 
the greatest climatic response across the United States. 

There appears to have been no growth response of Colorado Douglas-fir to Krakatau (6°S), 
unless there was a weak effect delayed until 1885. The eruption of Santa Maria (15°N) occurred 
in 1902, a year of intense drought on the Colorado Plateau (Cleaveland 1983, Appendix 1). The 
growing season of 1903 had adequate precipitation, and growth was slightly above average, but 
1904 was very dry resulting in low growth (Figure 2). The lack of adequate precipitation would 


certainly curtail the possible response of tree growth to volcanic cooling. Climatic effects from 
the eruptions of Katmai (58°N) in 1912 and Agung (8°S) in 1963 are also not discernible in these 
chronologies (Figure 2). The growth anomaly at 1816 is clearly the largest apparent in these data, 
and is probably the only one that can definitely be attributed to a known major volcanic eruption. 
However, there are other pronounced increases in growth that may be associated with 
undocumented volcanic activity. The possible detection of other volcanically created climatic 
effects in Colorado Plateau tree growth deserves further study. 

z 4 
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o me a Ci a nea) ee) am ret (oa Sao | (Le a aig | (omer Cea aD eae) Fee ee a Uf A Rg 
1487 1500 1550 1600 1650 1700 1750 1800 1850 1900 1950 1978 

Figure 2: Plots of tree-ring chronology index series from southwestern Colorado. A = average latewood 
width from Bobcat, Ditch, and Spruce Canyon sites; B = Ditch Canyon latewood width; C = 
Spruce Canyon latewood width; D = Bobcat Canyon latewood width; E = Bobcat Canyon ring 
width; F = Spruce Canyon latewood density (with variability multiplied by 3.0 to increase 
variability relative to the other variables). The numbers above each chronology are the sample 
depth at that point. 

The atmospheric mechanisms that may create regional climatic anomalies in response to volcanic 
influences are not well understood. Atmospheric and oceanic conditions at the time an eruption 
occurs may determine climatic response. It is believed, for example, that sea-surface temperatures 
and the El Nifio-Southern Oscillation phenomenon have mediated climatic effects of several recent 
volcanic eruptions (Angell and Korshover 1985; Angell 1988). 


The use of moisture-stressed conifers to investigate spatial patterns of historic volcanic eruption 
effects on climate may partially compensate for the limited distribution of instrumental climatic 
records prior to the twentieth century. Latewood width is particularly sensitive to the growing 
season moisture budget. The 1816 annual rings investigated in this report have the largest amount 
of latewood growth on the Colorado Plateau in 491 years. The pattern of greatly increased ring 
width, latewood width, and latewood density in these Colorado Plateau conifers from 1815 to 
1817 indicates a reduction of growing season moisture stress unique in the last five centuries. The 


climatic effect that apparently reached a maximum over the Colorado Plateau in 1816 probably 
began shortly after the April 1815 eruption of Tambora, and persisted into the growing season 
of 1817. The cause of this extraordinary growth anomaly was probably a reduction of mid- 
summer evapotranspiration demand, which appears to have extended the growing season, 
resulting in extremely large latewood growth. No effects of other known large volcanic eruptions 
were detected in these tree-ring chronologies. The receptivity of the general circulation to 
volcanic forcing may partly explain the apparently strong climate-tree growth response on the 
Colorado Plateau to the Tambora eruption, and the absence of a large growth response to other 
major eruptions during the historic period. Effects of eruptions on regional temperature and tree 
growth might be masked by existing regional climatic conditions such as drought, or by other 
climatic-forcing mechanisms such as sea-surface temperatures and/or the phase of the El Nifio- 
Southern Oscillation. This study has focused on annual ring data from a small set of chronologies 
in a small part of the Colorado Plateau, but it demonstrates a potential application of tree-ring 
data to the analysis of volcanic effects on climate. 


Thanks are due David W. Stahle, University of Arkansas Tree-Ring Laboratory, for suggested 
improvements to the manuscript, and to Thomas Harlan, University of Arizona Laboratory of 
Tree-Ring Research, who assisted in the collection and dating of the samples. Part of the data 
comes from my doctoral dissertation, which was supported by the Laboratory of Tree-Ring 
Research and the Department of Geosciences, University of Arizona. Additional support from 
the National Science Foundation, Climate Dynamics Program (grant ATM-8612343) is also 


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eruptions. Journal of Geophysical Research 93D:3697-3704. 

Angell, J.K. and J. Korshover. 1985. Surface temperature changes following the six major 
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Bradley, R.S. 1988. The explosive volcanic eruption signal in northern hemisphere continental 
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Bryson, R.A. and B.H. Goodman. 1980. Volcanic activity and climatic changes. Science 
207: 1041-1044. 

Cleaveland, M.K. 1983. X-ray densitometric measurement of climatic influence on the intra- 
annual characteristics of southwestern semiarid site conifer tree rings. Ph.D. dissertation. 

University of Arizona, Tucson. 177 pp. 

. 1986. Climatic response of densitometric properties in semiarid site tree rings. Tree-Ring 
Bulletin 46:13-29. 


. 1988. Corrigendum to: Climatic response of densitometric properties in semiarid site 
conifer tree rings. Tree-Ring Bulletin 48:41-47. 

Cook, E.R. and K. Peters. 1981. The smoothing spline: a new approach to standardizing forest 
interior tree-ring width series for dendroclimatic studies. Tree-Ring Bulletin 41:45-53. 

Drew, L.G. (ed.). 1976. Tree-ring Chronologies for Dendroclimatic Analysis. University of 
Arizona Laboratory of Tree-Ring Research, Tucson. 64 pp. 

Fritts, H.C. 1976. Tree Rings and Climate. Academic Press, London. 567 pp. 

Fritts, H.C., D.G. Smith and M.A. Stokes. 1965. The biological model for paleoclimatic — 
interpretation of Mesa Verde tree-ring series. American Antiquity 31:101-121. 

Fritts, H.C. and J.M. Lough. 1985. An estimate of average annual temperature variations for 
North America, 1602 to 1961. Climatic Change 7:203-224. 

Gilliland, R.L. and S.H. Schneider. 1984. Volcanic, CO, and solar forcing of northern and 
southern hemisphere surface air temperatures. Nature 310:38-41. 

Hammer, C.U., H.B. Clausen and W. Dansgaard. 1980. Greenland Ice Sheet evidence of post- 
glacial volcanism and its climatic impact. Nature 288:230-235. 

Harshvardhan and R.D. Cess. 1976. Stratospheric aerosols: effect upon atmospheric temperature 
and global climate. Tellus 28:1-9. 

Horstmeyer, S.L. 1989. In search of Cincinatti’s weather. Weatherwise 42:320-327. 

Kelly, P.M. and C.B. Sear. 1984. Climatic impact of explosive volcanic eruptions. Nature 

Lamb, H.H. 1970. Volcanic dust in the atmosphere; with a chronology and assessment of its 
meteorological significance. Philosophical Transactions of the Royal Society of London 

Lough, J.M. and H.C. Fritts. 1987. An assessment of the possible effects of volcanic eruptions 
on North American climate using tree-ring data, 1602 to 1900 A.D. Climatic Change 

Newhall, C.G. and S. Self. 1982. The volcanic explosivity index (VEI): an estimate of explosive 
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Panshin, A.J. and C. de Zeeuw. 1970. Textbook of Wood Technology, Vol. 1, Third Edition. 
McGraw-Hill Inc., New York. 705 pp. 

Parker, M.L., R.D. Bruce and L.A. Jozsa. 1980. X-ray densitometry of wood at the W.F.P.L. 
Forintek Canada Corporation Western Laboratory, Technical Report No. 10:1-18. 


Pollack, J.B., O.B. Toon, C. Sagan, A. Summers, B. Baldwin and W. Van Camp. 1976. 
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Rampino, M.R. and S. Self. 1984. Sulphur-rich volcanic eruptions and stratospheric aerosols. 
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Shutts, G.J. and J.S.A. Green. 1978. Mechanisms and models of climatic change. Nature 

Stockton, C.W. and H.C. Fritts. 1971. Conditional probability of occurrence for variations in 
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Paleoclimate Analysis and Modelling. A.D. Hecht (ed.). John Wiley and Sons, New 
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Stokes, M.A. and T.L. Smiley. 1968. An Introduction to Tree-Ring Dating. University of 
Chicago Press, Chicago. 73 pp. 

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Island. 177 pp. 

Stothers, R.B. 1984. The great Tambora eruption in 1815 and its aftermath. Science 224:1191- 

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1816 in Perspective: the View from the Northeastern United States 

William R. Baron! 

The year 1816 is remembered in the northeastern United States as one of the harshest, coldest 
years on record, and continues even to the present to be one of the most widely known folk- 
climate episodes of the region. A study of the period 1790-1839 helps to place 1816 in its 
climatological context. New evidence supports the case that 1816 had a particularly cold and dry 
growing season but was by no means the coldest or driest year of the period. Several other years 
during the second and fourth decades of the century climatically rivaled the abnormal conditions 
of the "year with no summer". 1816’s claim to fame rests on the severe impact that the cold and 
drought had on the area’s then extensive agricultural operations. For this reason 1816 came to 
represent the abnormal climatic conditions of not only a single year but also most of the second 
decade of the nineteenth century. 


The year with no summer, 1816, is one of the best known folk-weather occurrences of the 
northeastern United States. During the nineteenth century, "eighteen-hundred-and-froze-to-death" 
was a Subject for many newspaper articles, autobiographical reminiscences, and local histories 
(Mussey and Vigilante 1948). Even now, some 170 years after that frosty summer, it continues 
to be a topic of great popular interest, still commanding feature articles in the region’s 
newspapers and periodicals (Fichter 1971; Leach 1974; Reichmann 1978; Parsons 1980). 

1816 also has not escaped the attention of scientific investigators, and has been a subject of 
considerable debate since the early nineteenth century (Skeen 1981). During the twentieth 
century, research appearing on this topic included that by W.I. Milham (1924), J.B. Hoyt (1958), 
H.E. Landsberg and J.M. Albert (1974), H. and E. Stommel (1979, 1983), and R.B. Stothers 
(1984), and centred on the issue of what factors contributed to 1816’s abnormal summer. Most 
researchers concluded that the great Tambora eruption of 1815 and low sunspot activity were the 
major factors involved; although a minority, including Landsberg, have questioned the influence 
of volcanic dust in the atmosphere as a major contributor. Of late, historians have begun to assess 
the climatic impact of the 1810s on society in both Europe and North America (Post 1977; Skeen 
1981). Our continuing interest in and fascination with 1816 finally led to the international 
conference that produced the papers included in this volume. 

The purpose of this study is to present new and additional evidence for the northeastern United 
States covering the period 1790 through 1839, in order to help place 1816 in its proper 
climatological context. Data presented include instrument readings for temperature, precipitation, 
and wind direction. Additional reconstructions for snowfall, seasonal precipitation, cloud cover, 
growing-season length, thunderstorm frequency, river freeze-up and ice-out, and phenology 
records (based on the analysis of qualitative materials such as diaries), weather journals and 
newspaper reports, also are discussed. However, before proceeding, I will present a 

'! Historical Climate Records Office, Center for Colorado Plateau Studies, Northern Arizona University, Box 5613, 

Flagstaff, Arizona 86011, U.S.A. 


reconstruction of the weather history of 1816 based on the observations of 74 diarists from 
northeastern United States. 

1816, "The Year With No Summer" 

Appropriately enough, 1816 began, at least in Phillipstown, Massachusetts, with enough snow 
on the ground for sleighing. All over New England, January was a snowy, stormy month until 
the very end when a sudden thaw caused localized flooding such as the one reported by Isaiah 
Thomas at Worcester, Massachusetts on 23 January where some mill dams were carried off and 
some items stored in a warehouse were destroyed. 

According to among others, Leonard Hill of East Bridgewater, Massachusetts, February was a 
mild and pleasant month with only three snows reported. By the beginning of March there was 
little deep snow anywhere with the exception of most of northern New England. Early March was 
clear and cold, and was followed by a series of three snow storms around mid-month that 
produced a few days of sleighing but soon melted. On 28 and 30 March, warm air returned 
producing thunder and lightning as reported by Elijah Kellogg at Portland, Maine and Thomas 
at Worcester. 

April quickly turned cold again with frequent frosts and some snow. However, by 14 April, there 
was little snow left at Hallowell, Maine. By 19 April, Alexander Miller of Wallingford, Vermont 
had begun to plough his fields; Stephen Longfellow of Gorham, Maine was already planting 
wheat; and Theodore Lincoln of Dennysville (in far downeast Maine) was reporting ice-out on 
the local streams - a sure sign of coming spring. At the end of the month, Joshua Lane of 
Sanbornton, New Hampshire already was reporting the start of a drought that would later plague 
all of northern New England. 

In early May, farmers throughout the region completed planting their major crop, corn (maize). 
By mid-month the weather had become "backward" again with a “heavy black frost" that froze 
the ground to at least one-half inch (1.3 cm) reported on 15 May. as far south as Trenton, New 
Jersey. Miller, at Wallingford, Vermont, reported snow on 14, 17 and 29 May while Lane, over 
at Sanbornton, saw a large frost on 24 May, and ended the month with further complaints about 
the continuing drought. B.F. Robbins, visiting Concord, New Hampshire noted that May ended 
with two days of "remarkable cold" that froze the ground "to near an inch". 

June is the month most remembered for its outbreak of cold weather. On 4 June, there were 
frosts at Wallingford, Vermont and Norfolk, Connecticut. By 5 June the cold front was reported 
over most of northern New England. On 6 June, snow was reported at Albany, New York and 
Dennysville, Maine, and there were killing frosts at Fairfield, Connecticut. 

7 June brought reports of severe killing frosts from across the region, and as far south as 
Trenton, New Jersey. 

Typical of comments by diarists concerning this day are those by George W. Featherstonhaugh 
of Albany, New York, who wrote that the frost killed most of the fruit, as many apple trees were 
then just finishing blossoming. Leaves on most of the trees were "blasted" by the cold. Corn and 
vegetable crops were injured. He also feared many of the sheep that had just been sheared might 
die of cold. 

Cold weather continued through the night of 10 June. By the end of the month most observers 
were reporting the return of warm weather, but by then most crops were either killed or 


"backward" and stunted in their growth. In northern New England, those crops that survived the 
frosts were hit by what was now a very serious drought, greatly reducing production of one of 
the area’s primary crops, hay. 

In early July there was another outbreak of cold weather in northern New England. On 5 July, 
at Gorham, Maine, there was a very hard frost. Benjamin Kimball of Concord, New Hampshire 
and Thomas Robbins of Norfolk, Connecticut reported hard frost on 7 July. There was frost on 
8 July at Portland, Maine and on the following day at Sanbornton, New Hampshire. Thereafter 
the cold held off for the remainder of the month. Throughout the entire month dry conditions, 
generally reported earlier in northern areas, persisted. 

Frosts returned on the morning of 21 August, being reported at York and Portland, Maine and ~ 
Wallingford, Vermont. By 22 August hard frosts were noted all over the region and as far south 
as Trenton where buckwheat crops were killed. Thomas, at Worcester, Massachusetts, reported 
that these frosts "cut off Indian corn in many places", while others such as Hill at East 
Bridgewater, Massachusetts observed that frosts did little or no damage. 

The frosts continued into September. In northern New England there were frosts on 10 and 11 
September and throughout New England during 25-27 September. On 28 September, there was 
a killing frost throughout the region extending as far south as Trenton. It killed any vegetation 
that had somehow survived to that date. The drought in northern New England was finally broken 
by rains in the last week of the month. 

The remainder of autumn was very mild with very few snowfalls or storms. December was also 
mild, until the last 10 days or so, when it turned cold enough to freeze the harbour at Beverly, 
Massachusetts. The year ended with enough snow on the ground at Phillipstown, Massachusetts 
to use a sleigh. 

The place of 1816 in the memory of the regional population has been summed up well by the 
historian H.F. Wilson when he wrote that, in 1816, farmers experienced an "almost total failure" 
of major crops. There was a fair yield of winter grain, but other crops such as corn and hay 
failed leading to the loss of many sheep and cattle for lack of feed during the following winter. 
As a result 1816 has come down to us as the "cold year", "the famine year" and "eighteen 
hundred and froze to death". 

Of great interest to climatologists and historians alike is the fact that 1816 was not the only 
difficult, abnormal year of the second decade of the nineteenth century. Based on statistical 
analysis of climatological data, other years might justifiably claim a portion of 1816’s notoriety. 
An analysis of the 50 years surrounding 1816 serves to locate some of these years and to place 
the "year with no summer" in its climatological context. 

Databases and Methodologies 

The analysis that follows is based upon records assembled from several databases. The first of 
these is a set of eight yearly mean temperature records comprised of instrument readings for 
periods of 26 years or more, that overlap 1816 by at least three years and that, in composite, 
cover the 50 years from 1790 through 1839. From north to south these records include: Castine, 
Maine, 1809-39 (Baron et al. 1980); Brunswick, Maine, 1807-36 (Cleaveland 1867); Salem, 
Massachusetts, 1790-1829 (Holyoke 1833); New Bedford, Massachusetts, 1813-39 (Rodman 


1905); Williamstown, Massachusetts, 1811-38 (Milham 1950); New Haven, Connecticut, 1790- 
1839 (Landsberg 1949); New Brunswick, New Jersey, 1790-1839 (Reiss et al. 1980); and 
Philadelphia, Pennsylvania, 1790-1839 (Landsberg et al. 1968). All but one of these stations, 
Williamstown, are located in the present coastal climatic zone as computed by the United States 
National Oceanic and Atmospheric Administration. Williamstown’s inland situation makes it more 
vulnerable to outbreaks of Arctic cold coming from the northern interior of the continent. 

Instrumental records of precipitation for the period are more limited than those for temperature. 
The three best include those for New Bedford, 1814-39 (Rodman 1905); New Haven, 1804-21 
(Kirk 1939); and Philadelphia, 1790-1839 (Landsberg et al. 1968). No long precipitation records 
for inland locations are available. 

To assure the representativeness of these long-run temperature and precipitation records, and to 
enhance the record density and distribution throughout the region, a database of short-duration 
instrumental records was assembled. The location of these records is shown in Figure 1. The 
number of records available increases, by decade, in a steady progression from the 19 available 
in the 1790s to 57 for the 1830s. 

Yet another database (qualitative materials from diaries, journals, newspapers, and local histories 
- many available only in manuscript form) was compiled to supplement the instrumental records. 
At that time, only qualitative materials for New England were sufficiently organized for 
inclusion. The database is comprised of 174 sources representing 55 of New England’s 67 
counties. Coastal and intermediate interior locations are well represented whereas data for some 
upland and western interior counties are missing. From this database, frequency counts for days 
with precipitation, fair skies, thunder and lightning storms, westerly winds, and snowfalls, as well 
as the yearly dates for spring and autumn killing frosts, length of snowfall seasons, dates of 
apple-tree blossoming and lengths of droughts were compiled. The methodology used to 
reconstruct these various records is explained in Baron (1988, 1989), Baron and Gordon (1985) 
and Baron et al. (1984). 

Compilations of killing-frost reports were further refined by computing the year and day of each 
frost and calculating the length of the growing season. To assess the possible impact of these 
frosts on a major crop, information from an agricultural database made up of over 60 farm 
journals was used to provide the mean dates for the planting and harvesting of Indian corn. 

Analysis of Records 

As can be seen in part from Figure 2, moving from north to south, 1816’s yearly mean 
temperature is not the coldest for the period. In Maine it was only the fourth lowest; while farther 
south at Salem, New Haven, and New Bedford it was third or second lowest for the record. Even 
farther south at New Brunswick and Philadelphia, 1816 was close to the mean for the period 1790 
through 1839. Farther inland at Williamstown, 1816 was the seventh coldest year (Figure 3). In 
New England the years noted to be as cold or colder than 1816 include: 1790, 1812, 1817, 1818, 
1823, 1835, 1836 and 1837. 1836 and 1812 appeared most often as the two coldest years. South 
of New England, 1836 and 1817 were the coldest years. 


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Figure 1: Northeastern United States. Instrumental record locations, 1790-1840. 






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Castine, Maine (A); Salem, Massachusetts (B); and New Brunswick, New Jersey (C); annual 

mean temperatures, 1790-1840. 


5 Year Running Mean 

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——t— Annual Mean Temperature 




Figure 2 



——&— Annual Mean Temperature Year ow —5-Year Running Mean 





Figure 3: Williamstown, Massachusetts. Yearly mean temperatures, 1811-38. 

Taking all temperature records together for the 50-year span, apparently the years 1790 through 
1798 were either normal or slightly cooler than normal. 1799 through 1805 was warmer than 
normal. A period of variability between 1806 and 1811 followed. Starting in 1812 (with some 
variation until 1818) it was much cooler than normal. This cool period was followed by a series 
of variable years from 1819 through 1823. From 1824 through 1830 it was somewhat warmer 
than normal, and from 1831 through 1837 it was again very cool. The last two years of the 
decade show a warming trend that extended into the 1840s. 

Figure 4 shows an analysis of seasonal mean temperatures for Brunswick and New Haven. Winter 
mean temperatures were calculated from December, January and February monthly means; spring 
means from March, April and May; summer means from June, July and August; and autumn 
means from September, October and November. The Brunswick record shows the winter of 1816 
was very cold, whereas the summer was cool, the spring about average and the autumn a little 
warmer than average. Summer mean temperatures in 1812 were as low as those for 1816. New 
Haven presents a somewhat different picture with a mild winter, average autumn and cool spring 
and summer. As far as New Haven is concerned, the summer mean temperatures for 1812 and 
1817 were far lower than that for 1816. 

Figure 5, a daily mean temperature record for 1816 kept at Brunswick, Maine shows why this 
year is so well remembered. The outbreaks of cold in much of May, early June, early July, late 
August and late September tell much of the story. These cold periods, all well below the monthly 
mean temperature for the entire record (1807-36), doomed many a farmer’s crops to failure. The 
key to 1816’s infamy lies in the extreme shortness of its growing season - the primary reason 
why it, and not 1812 or 1836, has gone down in the regional weather lore as "the year with no 



A 25 



B 30 


———— Summer —o— Fall 



—e— Spring 

Brunswick, Maine (A) and New Haven, Connecticut (B); winter, spring, summer, and autumn 

season mean temperatures, 1790-1840. 

——— Winter 

Figure 4 

Daily Mean Year Days 

Monthly Means for Record 

0 30 60 90 120 150 180 210 240 270 300 330 360 

Jan. Feb. March Apr. May June July Aug. Sept. Oct. Nov. Dec. 



eT A) 



Figure 5: Daily mean temperatures at Brunswick, Maine for 1816. 

The plots of growing-season lengths for eastern Massachusetts, southern New Hampshire and 
southern Maine (Figure 6) leave one unmistakable impression - 1816, by far, has the shortest 
growing season. Other particularly short growing seasons occurred in 1808, 1824, 1829, 1834 
and 1836. With the exception of 1816 and 1836, a number of these short seasons can probably 
be attributed to one-night radiational cooling under clear skies during either spring or autumn, 
and not to prolonged outbreaks of cold weather; otherwise these years would have appeared in 
our lists of cool yearly and seasonal temperatures. Of course clear skies in combination with cold 
fronts also contributed to some frosts during 1816, as that year has one of the highest percentages 
of days with fair skies (Figure 7). 

Figure 8, showing spring and autumn killing-frost dates in combination with corn-planting and 
maturation dates, further illustrates the importance of growing-season data in understanding the 
notoriety of 1816. For eastern Massachusetts, 1816 is the only year in which young corn was 
killed in the spring after it had sprouted and in which corn that survived replanting was killed in 
the autumn, before it could reach maturity. Under these circumstances, it is safe to assume that 
in most places in New England corn crops were an almost total failure. The story for 1816 is the 
same for New Hampshire and Maine. These reconstructions also show there were a number of 
years when corn crops were hit by late spring or early autumn frosts. Particularly difficult periods 
include: 1793-96, 1812-17, and 1835-36. 





C 240 


Figure 6: Growing-season lengths for southern Maine (A), 

_ - | 

| ay 

~~ DO = 
ma Dm oO 
Sf & 








z co 


ean of Growing Season 

Massachusetts (C), 1790-1840. 









fo>) — w ~ a = wo ~ fo?) _- 
cS" N N N N oO oO oD oO SSE 
co co co co co co foe] co co foe) 
a = = ==. = 222 
\\ Bilt 
| | 
a — oO w ™~ for) = oO wo ~ a> - 
ts N N N N N oO oO oO (se) oO wr 
co foo] foo] co co co f°) co co co co foe) 
S28 2353 2353 32 2282 
PX s = (\ 
~ || \ / 
} ¥ 
a = oO wW ~™ fo) S Oo wo ™~ fo?) — 
_ N N N N N oO oO oO Oo oO aT 
co Loo] ce co foo} t- 2] co foo} co co co foe) 
Ses eaa a 8B S253 322 
ss 5 Year Running Mean 

southern New Hampshire (B) and eastern 


Opn nnnnnnnncnanstnnnnnnnnnnnsnnantunnnnnnnnnnanaaannanannnnnnannannnnnntnnnannannnnantnnaniannnnnnansatanntanuntnnnnnnnnniaiannnnntaniniinnananantnninnnininssnnnannansnnansnsnsananianenisnnansnn 




































Percentage of Fair Days per Year 



Spring Year Hi Summer Fall 

Figure 7: Percentage of days with fair skies for eastern Massachusetts, 1790-1840. 

Reconstructions indicating the onset of spring-like weather, such as dates of last snowfalls and 
blossoming dates of various fruit trees such as apples (Figure 9), show that 1816’s spring was 
cooler and more unsettled than normal. However, it was far more satisfactory for agricultural 
pursuits than 1812, 1832, or 1836 through 1838, when conditions were extremely "backward". 

While the major part of the 1816 story lies in its growing-season record, there are several other 
types of records in which 1816’s position is worthy of mention. The first of these is the 
precipitation record (Figure 10). 1816 was a year of about average precipitation, with the 
exception of the summer, which was particularly droughty. Reconstructions concerning the period 
and intensity of agriculturally-defined droughts show that the magnitude of dry conditions 
increased significantly the farther north within the region one looks. For the 50 years from 1790 
through 1840, the periods from 1791-1806 and 1813 to 1820 saw numerous growing-season 
droughts. After 1820 the number of reported droughts decreased markedly. This apparent 
increase in precipitation also can be seen in Figure 11, illustrating the mean number of days per 
year with precipitation for southern New England. 




AeA fl 



ii vill 
WA ii 





LAU ll rt 
Mi i hil iit i Lae 

ital il 









vie i 

Year Day 

I a ne rere ne rnrerer errr rece rernererenerined 







Wad TTT 
it Ly i ‘ Hi 



° wevererer: 







—s— FallFrost 



120 Day Grow season 




Year Day 
Year Day 

—e—_ Corn Plant Date 






Killing frost and corn plant/harvest dates for southern Maine (A), southern New Hampshire 

(B), and eastern Massachusetts (C), 1790-1840. 


a OZ) 
co co 
- = 




—t— Spring frost 

Figure 8 








































Figure 9: Apple-blossom dates for northern (A) and southern (B) New England, 1790-1840. 


FJ Fall 



B.S Se sei 

ds cee 
a Lea! — 
a i 
= = 4 
Basttttted NNN NNN A ENE RS a. 
WN ard 628 : en 
| | | aa x : ieee 
— 3 ea i 
| ____ N  8 a ima EY 
JE sca : aes ses ee a 
yal E fae cae me 
RONAN 8281 2 eae oo 
ci Pasa aie as aN 
2281 g Seas MRO aa PY 
oe 5. ne RS Sa 
- 5 igen re 
wn Sasi 
____ Ny a : oe 
ANN 18h se ead ae 
2 | 3 a eee here 
© 2 NNW eee 9181 > Heil cel ea ea 
a = 5 Gos ecm —_ 
eZ (eee es. 
- eh ; ee 
é [a ol el [77 
gg 2 Re oe ae 
= = = - = _ - 

uojyey|djoesg YM seq 


Figure 11: Mean number of days per year with precipitation for southern New England, 1790-1840. 

PAAR Rares 
UREA a a a A 

jjeymous ujIM sheq jo Jsaquiny 


E83 Spring (Mar-May) 
E] Winter (Dec-Feb) 

HE Fall (Oct-Nov) 

|Je}MOUS Y}IM SAep jo Jaquiny 


Number of days with snowfall in southern Maine (A) and eastern Massachusetts (B), 1790- 


Figure 12 

Snowfall records show that during the winter of 1815-16 there were relatively few days when it 
snowed, especially in northern New England (Figure 12), but the New Bedford seasonal 
precipitation record shows average or slightly above average precipitation for that winter. 
However, under no circumstances can 1816 be viewed as a particularly snowy winter. Available 
records led me to believe that in northern New England, during 1816, winter conditions were 
drier and colder than normal; while farther south in Massachusetts the season was warmer and 
wetter. Among the years with the greatest number of snowfalls were: 1792, 1804, 1805, 1806, 
and 1818. Those years with the least number included: 1828, 1829 and 1834. 

Reconstructions bearing evidence concerning the storminess of the period [e.g., the percentage 
of days with westerly fair-weather winds (Figure 13) and the number of days in which thunder 
and lightning storms occurred (Figure 14)], show that 1816 was rather stormy. The frequency 
of westerly "fair weather bearing" winds was somewhat below the record mean; while days with 
thunder and lightning storms numbered close to the record mean. Evidently there was a decrease 
in storm activity during the 1790s. In the early 1800s, there was a small increase followed by 
another decrease late in the decade. From 1809 through 1822, there was considerable year-to-year 
variability but an overall increase in storminess. During the late 1820s and all of the 1830s, there 
was a general decrease in storm frequency. 

70. f i F i : i 

‘ Al) 
ill il 



% Westerly Winds 



Wo [ti 


ic a 
~ co 
- _- 

—t— % Westerly Winds by Year Year 5 Year Running Mean 

Figure 13: Westerly winds per year over southern New England, 1790-1839. 




RS Shhh AAA, VQAa&G@&K ws 8e8l 
= Leet 

Sy SOON resi 
REGAN & zest 
SS Mih&nnAHnR] AAA VVSSVVTsssss Oe8l 
SE : SS y28l 

: c28l 

NRE 9181 

Se SSS Sts8l 
QRSE_iinn hhh] e181 

S55 1081 
Sas 008! 

SY Z6Lt 




wo o wo °o wo o 
N N 


Bujuyy8)q 9g sepuny, ym sheqg jo sequinn 


Number of days per year with thunder and lightning storms for eastern Massachusetts, 1790- 


Figure 14 


1790 through 1839 featured two abnormally cold periods (1812 through 1818 and 1832 through 
1838) and two warmer, relatively more stable periods (1799 through 1810 and 1819 through 

1830). The warmth and stability of the latter two decades, compared with cold and relative 
storminess of the 1810s, heightened peoples’ awareness of the contrast between the two climatic 

Especially in northern New England, where considerable farming took place on 

climatically-marginal lands, the cold years brought disaster. To make matters worse, the swing 
from warm to cold in the 1810s coincided with an increase in economic competition from the 


midwestern United States and central Canada. The additional stress of crop shortfalls due to 

shortened growing seasons forced many farmers to leave New England for what they believed 

were more hospitable climates to the west (Smith et al 1981). 

1816 was only one of several abnormal years that occurred during 1790 through 1839. When 

viewed from this perspective, 1816’s abnormality pales. Why then is 1816 so well remembered 

s fabled years of climatic 


while 1812 or 1836 are assigned to the second rank of the region 



The answer lies not in our careful compilation of climatological records (for statistically 1816 
does not measure up) but in the nature of 1816’s abnormality and the impact of its greatly 
shortened growing season on New Englanders’ capability to raise food. A harsh, snowy winter 
or severe spring flood have a great impact on certain segments of society, but a series of killing 
frosts accompanied by a severe drought (especially in northern New England) hit nearly the entire 
society by forcing up food prices for the rich and by reducing the available larder for the poor. 
For the average New Englander, particularly the farmer, 1816 was the worst year of a series of 
bad years. As time passed, 1816, in New England’s folk memory, came to stand for the 1810s 
as a whole. This idea has been passed down from generation to generation as the story of "the 
year with no summer". 

Acknowledgements and a Note on the Availability of Climatic Data 

I thank the staff of the Historical Climate Records Office for their assistance. This research was 
partially supported through Northern Arizona University’s Organized Research Fund; I thank 
particularly Henry O. Hooper, Associate Vice President for Academic Affairs, Research and 
Graduate Studies for his support and assistance. 

All databases discussed here are kept in the Historical Climate Records Office, part of the Center 
for Colorado Plateau Studies at Northern Arizona University, Flagstaff. The Office has on file 
and computer disc a large number of United States records collected for the seventeenth through 
nineteenth centuries. There are particularly strong record groups for the northeastern United 
States and the Colorado Plateau. The Office was founded by the author with the intention of 
making these climatic materials available to other researchers. Record collection was done by the 
members of the now disbanded Northeast Environmental Research Group centred at the 
University of Maine and, after 1985, by the staff of the Historical Climate Records Office. 
Record collection undertaken through 1985 was supported by grants from the National Science 
Foundation and the Northeast Regional Experiment Stations to the University of Maine. A listing 
of available records may be obtained by writing me. 


Baron, W.R. 1988. Historical climates of the northeastern United States: seventeenth through 
nineteenth centuries. Jn: Holocene Human Ecology in Northeastern North America. G.P. 
Nicholas (ed.). Plenum Press, New York. pp. 29-46. 

. 1989. Retrieving climate history: a bibliographical essay. Agricultural History 63. (in 

Baron, W.R., D.C. Smith, H.W. Borns, Jr., J. Fastook and A.E. Bridges. 1980. Long-Time 
Series Temperature and Precipitation Records for Maine, 1808-1978. In: Life Sciences and 
Agriculture Experiment Station Bulletin 771. University of Maine Press, Orono, Maine. 
p: 97. 

Baron, W.R., G.A. Gordon, H.W. Borns, Jr. and D.C. Smith. 1984. Frost-free season record 

reconstruction for eastern Massachusetts 1733-1980. Journal of Climate and Applied 
Meteorology 23:317-319. 


Baron, W.R. and G.A. Gordon. 1985. A reconstruction of New England climate using historical | 
materials 1620-1980. In: Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 

Cleaveland, P. 1867. Results of meteorological observations made at Brunswick, Maine, between 
1807 and 1859, reduced and discussed by C.A. Schott. Smithsonian Contributions to 
Knowledge 16:2-25. 

Fichter, G.S. 1971. Eighteen-hundred-and-froze-to-death: snowfilled summer of 1816. Science 
Digest 69(2):62-66. 

Holyoke, E.A. 1883. A meteorological journal from the year 1786 to the year 1829, inclusive. 
Memoirs of the American Academy of Arts and Sciences, New Series 1:107-216. 

Hoyt, J.B. 1958. The cold summer of 1816. Annals of the Association of American Geographers 

Kirk, J.M. 1939. The Weather and Climate of Connecticut. State Geological and Natural History 
Survey Bulletin 61. Connecticut Geological and Natural History Survey, Hartford. 242 pp. 

Landsberg, H.E. 1949. Climatic trends in the series of temperature observations at New Haven, 
Connecticut. Geografiska Annaler 1(2):125-132. 

Landsberg, H.E., C.S. Yu and L. Huang. 1968. Preliminary reconstruction of a long time series 
of climatic data for the eastern United States. Institute for Fluid Dynamics and Applied 
Mathematics, University of Maryland, Technical Note B14-571:1-42. 

Landsberg, H.E. and J.M. Albert. 1974. The summer of 1816 and volcanism. Weatherwise 

Leach, A. 1974. 1816 was a year for complaints. Barre-Montpelier Vermont, Times Argus, 
22 July. 

Milham, W.I. 1924. The year 1816 - the causes of abnormalities. Monthly Weather Review 

. 1950. Meteorology in Williams College. McClelland Press, Williamstown. 40 pp. 
Mussey, B. and S.L. Vigilante. 1948. "Eighteen-hundred-and-froze-to-death". The cold summer 
of 1816 and westward migration from New England. Bulletin of the New York Public 
Library 1948:56S. 
Parsons, M. 1980. An eyewitness report of 1816, the cold year. Bittersweet 4-1:39-41. 

Post, J.D. 1977. The Last Great Subsistence Crisis in the Western World. Johns Hopkins 
University Press, Baltimore. 240 pp. 

Reichmann, R. 1978. 1816 recalled as year without a summer. Providence, Rhode Island Sunday 
Journal, 17 September. 


Reiss, N.M., B.S. Groveman and C.M. Scott. 1980. Seasonal mean temperatures for New 
Brunswick, N.J. Bulletin of the New Jersey Academy of Science 1980:1-10. 

Rodman, T.R. 1905. Monthly, annual and average temperatures and precipitation at New 
Bedford, Mass., 1813-1904. Climate and Crops: New England Section, Annual Summary 

Skeen, C.E. 1981. The year without a summer: a historical view. Journal of the Early Republic 

Smith, D.C., H.W. Borns, Jr., W.R. Baron and A.E. Bridges. 1981. Climatic stress and Maine 
agriculture, 1785-1885. In: Climate and History. T.M.L. Wigley, M.J. Ingram and 
G. Farmer (eds.). Cambridge University Press, Cambridge. pp. 450-464. 

Stommel, H. and E. Stommel. 1979. The year without a summer. Scientific American 240:176- 

. 1983. Volcano Weather: The Story of 1816, The Year Without a Summer. Seven Seas 
Press, Newport. 177 pp. 

Stothers, R.B. 1984. The great Tambora eruption in 1815 and its aftermath. Science 

Wilson, C.M. 1970. The year without a summer. American History Illustrated 5(6):24-29. 

Wilson, H.F. 1967. The Hill Country of Northern New England: Its Social and Economic History 
1790-1939. AMS Press, New York. 455 pp. (Reprint of 1936 edition). 

Sources Directly Mentioned in 1816 Weather History 

Bascom, R.H. 1816. Manuscript diary for Phillipstown, Massachusetts. American Antiquarian 
Society, Worcester, Massachusetts. 

Dickerson, M. 1816. Manuscript diary for Morristown and Trenton, New Jersey. New Jersey 
Historical Society, Newark, New Jersey. 

Featherstonhaugh, G.W. 1816. Manuscript diary for Albany, New York. Albany Institute, 
Albany, New York. 

Hill, B.T. (Editor). 1910. The diary of Isaiah Thomas 1805-1828. American Antiquarian 
Society Transactions and Collections 10:251-263. 

Hill, L. 1869. Meteorological and Chronological Register. L. Hill, Plymouth, Massachusetts. 
pp. 61-66. 

Kellogg, E. 1816. Manuscript notes in Old Farmer’s Almanack for 1816. Maine Historical 
Society, Portland, Maine. 


Kimball, B. 1856. Extracts of a diary. In: The History of Concord, 1725-1853. N. Bouton (ed.). 
B.W. Sanborn, Concord, New Hampshire. p. 771. 

Lane, J. 1816. Manuscript diary for Sanbornton, New Hampshire. New Hampshire Historical 
Society, Concord, New Hampshire. 

Larcom, J. 1951. Diary of Jonathan Larcom of Beverly, Massachusetts. Essex Institute Historical 
Collections 87:65-95. 

Lincoln, T. 1816. Manuscript record book of weather reports. Maine Historical Society, 
Portland, Maine, 1: (not paginated). 

Longfellow, S. 1816. Manuscript diary for Gorham, Maine. Maine Historical Society, Portland, 

Miller, A. 1816. Manuscript diary for Wallingford, Vermont. Vermont Historical Society, 
Montpellier, Vermont. 

Robbins, B.F. 1816. Manuscript travel diary. Maine Historical Society, Portland, Maine. 

Robbins, T. 1986. Diary of Thomas Robbins, DD, 1796-1854. Beacon Press, Boston. pp. 236- 

Sewall, H. 1816. Manuscript diary for Hallowell, Maine. Maine State Archives, Augusta, Maine. 

Weare, J. 1912. The diary of Jeremiah Weare, York, Maine. New England Historical and 
Genealogical Register 66:77-79. 

Wheeler, W. 1930. Diary of William Wheeler. Yale University Press, New Haven. p. 221. 


Extension of Toronto Temperature Time-Series from 1840 to 1778 Using 
Various United States and Other Data 

R.B. Crowe! 


Daily maximum and minimum temperatures for the city of Toronto are archived from 1 March 
1840 to the present day. This lengthy time-series can be extended considerably by using standard 
differences in mean monthly temperatures between Toronto and some United States stations, the 
earliest of which dates from July 1778. In addition, there are considerable temperature data taken 
three times a day from another station in Toronto in the 1830s. These data were adjusted and 
monthly mean temperatures calculated. Mean July temperature for Toronto in 1816 is calculated 
to have been remarkably low (nearly 16°C). 


In December 1839, the British Government established a meteorological and magnetic 
observatory at Toronto, Ontario. Some sporadic observations began late in the month at Fort 
York (on the shore of Lake Ontario just west of the town, then called York). Fixed hourly 
observations of temperature commenced in the new year, but not until 1 March 1840 were 
regular daily maximum and minimum values recorded: on this date daily Archive readings begin. 
On 5 September 1840, the observation site was moved to the University of King’s College (now 
the University of Toronto) about 2 km north of the lakefront. Although a number of small 
changes in location occurred in later years, a relatively homogenous, high-quality data set extends 
from September 1840 to the present day. However, the rather large urban heat-island effect, 
which influences the records of all such large cities, is evident. 

The Toronto record from 1840 to the present comprises the longest continuous temperature time- 
series in the Canadian climatological Archive, and thus it is frequently used in analyses of long- 
term temperature trends. 

The purpose of this paper is to present a method of extrapolating the Toronto temperature time- 
series backwards from 1840, using various United States and other data. The earliest American 
data used in the analysis were taken in 1778. The significance of these data for 1816 is 

Sources of Early Climatic Data Used for Comparative Purposes 

Toronto Area 

The first fragmentary climatic data for Toronto were taken in the year 1801. These are contained 
in the Hodgins Papers in the Archives of Ontario, Toronto, dating from the late nineteenth 
century. The data are identical to those published in the Upper Canada Gazette at the time the 
observations were taken, so presumably Dr. Hodgins merely copied long-hand this original 

' Canadian Climate Centre, Environment Canada, 4905 Dufferin Street, Downsview, Ontario M3H 5T4, Canada. 


source. The data include the temperature and weather or sky condition at three fixed times a day. 
Similar data were published for a number of months in the same newspaper around 1820. Neither 
data set was lengthy enough to be used in this study. 

Later a longer, more useful data set was taken by Dade (1831-41) from January 1831 to April 
1841. Reverend Dade was the Headmaster of Upper Canada College, then situated close to the 
centre of town on the lake shore, not far from the later Fort York station. The thermometer was 
read two or three times a day at fixed times, but slight changes in reporting hours occurred 
during the decade, and occasionally only one observation was taken in a day. Only a few months 
were incomplete, except for an extended period from October 1838 to June 1839 when Dade 
returned to England for the winter. 

Periods of record for the various early Toronto stations used for comparative purposes are shown 
(Figure 1). Data from Fort York are combined in the Archive with those from the University 
station and identified as "Toronto" (no modifier), but is unofficially called "Toronto City". Data 
later than 1855 were not considered. 


1831-41 1eAo we LOWVILLE 

ANCASTER _1831-50 : 
® LEWISTON NM OCHECIER 1830-49 1826-63 827-49 
BUFFALO 1827-49 @ @1557.60 

2 1830-60 @ CORTLAND 


) @ 
1789-1795 MAVEN 


1817-767 | 


Figure 1: Early eastern North America climatic data. 


Remainder of Southern Ontario 

Data for Ancaster (about 65 km southwest of Toronto) were taken by Craigie (1835) from 
January 1835 to December 1845, and proved to be of limited use in the Toronto data extension. 
William Craigie was a surgeon who apparently tabulated daily maximum and minimum 
temperatures as well as fixed-hour readings. His thermometers "were in a northern exposure, five 
feet from the ground, and shaded from the effects of direct insolation and radiation from the 
sky". However, only newspaper tabulations of monthly means of the 9 a.m. and 9 p.m. 
observations survive. 

American Stations 

Mean monthly temperature data were abstracted from publications of the Smithsonian Institution 
(1927) for Albany, New Haven and New York City and the United States Weather Bureau (1932- 
37) for Albany, Baltimore and Rochester. Considerable monthly data were also available from 
grammar schools in New York State (Hough 1855, 1872). Data for Auburn, Buffalo, Cortland, 
Fairfield, Fredonia, Hamilton, Lewiston, Lowville, Oneida, Rochester (College) and Utica were 
used, other stations listed in the above publications having insufficient useful data. 

All stations actually used in the study are shown in Figure 2. Many months and years for most 
Stations were noted in the New York State grammar school records, and only the first and last 
years of data are shown. In all cases, data later than 1855 were not used. 

In the case of most of these early data, excepting Toronto (city), observations were taken with 
the thermometer attached to the north wall of a building. Recording maximum and minimum 
thermometers were not generally used. Monthly means were computed from two, three or more 
observations a day, and the time and number of daily observations frequently changed and were 
not consistent, either at a site or from one station to another. In addition, thermometers may not 
have been calibrated accurately or sufficiently shielded from insolation, and changes in exposure 
or siting may not have been recorded. 

Method of Estimation of Toronto Mean Temperatures 

Three distinct methods were employed in the calculation of Toronto mean temperatures due to 
significant differences in the form of the source data: monthly means at the American stations 
(calculated by a variety of methods depending on the station); daily data for Dade; and monthly 
means for 9 a.m. and 9 p.m. in the case of Ancaster. These were labelled Method "S", Method 
"D", and Method "A" (Figure 3). 

All three methods were employed whenever data permitted. In deriving the final Toronto 
estimates, however, Method "D" was chosen whenever Dade information was available. Thus, 
Method "S" was used up to December 1830, but Method "D" from January 1831 to February 
1840. For missing Dade months, Method "S" was substituted before 1835, but from this year on, 
a linear regression equation was used based on the 52 months when the calculated American data 
could be compared with both Dade and Ancaster calculations: 

T = -0.126 + 0.6045 T, + 0.4108 Ts, 
where T is the estimated Toronto monthly mean (°F), 

T, is the estimated Toronto mean using Method "A", 
and____T, is the estimated Toronto mean using Method "S". 


ee DADE (On Lakeshore) 


Figure 2: Early Toronto climatological records. 


| (Mar.)1840 —~» TORONTO 

pill 1831-1841 DADE 

1. 1778-1830: METHOD "Ss" 
2. 1831-1834: METHOD "D" 


3. 1835- (Feb.)1840: METHOD "D" 


Figure 3: Data used to estimate mean monthly temperatures at Toronto. 


Method "S" 

This method essentially involved calculating standard differences between various American 
stations and Toronto. For example, since August is normally 4.5°F (-15.3°C) cooler at Toronto 
than at Albany, it was assumed that all missing August means at Toronto for which data were 
available at Albany could be reasonably estimated by subtracting 4.5° from the Albany values, 
no matter whether the month was near normal or significantly below or above normal. 

Method "S" is outlined in Figure 4. There are five distinct steps: 


Difference calculations. Mean monthly temperature data for Toronto (city), 1840 (March- 
December) - 1870, and for all available American stations within about 400 miles (644 km) 
of Toronto having significant data before 1870 were tabulated. Sixteen distinct United States - 
stations were available. Most stations did not have data before 1820, but New Haven had data 
as early as 1778 (Figure 2). In order to facilitate comparison of data from all stations, for 
each individual month, differences were calculated for all possible stations pairs, for example, 
Toronto minus Rochester, Toronto minus Albany, Toronto minus New Haven, Rochester 
minus Albany, etc. 

Correction and deletion of bad data. The differences for each station pair were tabulated 
by month and year, and the overall monthly-mean differences calculated. Also, standard 
deviations of the mean monthly temperatures for each station were calculated. Then, for each 
station pair for each of the 12 months, an average standard deviation (s.d.) was calculated, 
and all differences greater or less than one standard deviation were identified. For example, 
the s.d. of the August means at Toronto is 1.8°, Albany 2.2°, for an average of 2.0°. The 
mean difference for the month, Toronto-Albany, is -4.5°, so that the 1 s.d. range is -2.5° 
to -6.5°. By identifying differences outside the 1 s.d. range, it was easy to spot unusual 
months or questionable data. For any month for which no station-pair differences lay outside 
the 1 s.d range, the data were assumed to be reasonably good. Otherwise, a subjective 
assessment was made by the areal plotting of the means and departures from normal. In a few 
cases, it was possible to correct a value when it was obvious that a typographical error of 
10°F (12.2°C) had been made in the printed source. Most questionable data, however, were 

Choosing of stations and periods for analysis. Following the corrections and the discarding 
of questionable months, a new data set for each of the 16 American stations was prepared. 
Toronto data were assumed to be "good", and the aim was to choose as many of the 16 
United States stations as possible for the 1840-70 period for comparative purposes. Many 
stations had missing or discarded data in the last half of this period, so it was necessary to 
use only 1840-55. Within this period, not enough data were available for the computation of 
reasonable monthly means at Auburn or Buffalo, and data for Rochester College were 
identical to, or varied by only a small constant from, the Rochester data and hence was 
suspect. The number of American stations useful for comparative purposes was. therefore 
reduced to 13. 

Standard difference calculations. For each of the 12 months, mean differences in monthly 
temperatures between Toronto and each of the 13 American stations were calculated. The 
calculations were based on all months, September 1840 to December 1855, inclusive. I 
considered that the early data for Toronto (March to August 1840), taken near the lake shore 
at Fort York, were not homogeneous with the later University site observations. Because of 


some missing months for most United States stations, standard differences were calculated 
from means based on 10 to 15 years in most cases, and which varied from month-to-month 
and from station-to-station. The standard differences between Baltimore and Toronto ranged 
between 10.0° and 13.5°, depending upon the month. Because of its great distance from 
Toronto, and the resulting high and variable differences in the monthly mean temperatures, 
I decided to eliminate Baltimore. Thus, only 12 stations were left for further analysis. 

5. Calculation of Toronto means. For each month, July 1778 to February 1840, an estimated 
mean temperature for Toronto was calculated separately based on each American station for 
which a mean monthly temperature was available. Thus, in June 1831, the mean monthly 
temperature at Albany was 72.8°F (22.7°C), and since the standard difference, Toronto- 
Albany (based on 1840-55) is -6.8°, the Toronto mean was estimated at 66.0°. Similarly, 
New Haven was 71.1° in the same month, and the standard difference, Toronto-New Haven, 
is -5.5°, so that Toronto’s mean was calculated at 65.6°. The overall Toronto mean was 
calculated as the unweighted average of all the individual estimates from the various 
American stations, which for any month would vary from one to 12. This method of 
estimating Toronto means was checked against the actual means from September 1840 to 
December 1855. Although some monthly errors were as great as 4°F, the standard deviations 
of the errors varied from 0.94 to 1.75 for individual months, averaging 1.17°. 

Method "D" 

In the case of Dade data, mean monthly temperatures had to be calculated from two or three 
hourly observations a day. Then, standard difference calculations were made by comparing Dade 
monthly means with those of several American stations. Toronto monthly means were also 
compared to the same United States stations for the period 1840-55. In this way, a first 
approximation was made of Toronto-Dade differences. The period of the Dade observations was 
somewhat cooler on average than that of the later Toronto observations. Consequently, a second 
approximation of the Toronto-Dade differences was made, allowing for mean temperature 
differences between the two periods. When Toronto means were calculated month-by-month using 
these second approximation differences, a comparison with the means obtained by using Method 
"S" showed a consistent low bias. Hence, a final approximation of Toronto-Dade differences, and 
therefore, calculation of Toronto means, was made to allow for this low bias. 

Method "D" is outlined in Figure 5. There are eleven distinct steps: 

1. Estimation of daily maximum and minimum temperatures. When Reverend Dade began 
observations on | January 1831, he took readings at 9 a.m., 3 p.m and 6 p.m. However, this 
pattern did not continue in later months and years. Sometimes the morning reading was at 
7 or 8 a.m., but the mid-day observation was usually taken at noon and the evening one at 
5 p.m. To complicate matters, while there was always a morning reading (excepting the 
months with partial days, which were excluded from analysis), sometimes there was in 
addition only a mid-day reading, sometimes only an evening one, sometimes both and 
sometimes neither. It was necessary, therefore, to estimate daily maximum and minimum 
readings. This was done by using mean hourly temperatures in comparison with mean daily 
maxima and minima for each month at Toronto’s Pearson International Airport (Atmospheric 
Environment Service 1978). Thus, a correction factor was calculated to be subtracted from 
the morning reading to estimate the daily minimum and to be added to the mid-day and 
evening observations to estimate the daily maximum. These applied to Pearson Airport, so 
corrections for Dade were computed by multiplying the Pearson corrections by the ratio of 



















Figure 4: Method "S". 


1 to 16 U.S. Stations 1778-1870 
and Toronto 1840-1870 

For each month all possible station pairs 
(GtnewAr= "Stn. Bo Stne A= Stn. ce 
Stale 13) = Suis (6, Geis. 

For each of the 12 months 
Sep. 1840 - Dec. 1855 mean 
differences in monthly 
temperatures between Toronto 
and each of 12 U.S. stations 

For each month July 1778 - 
Feb. 1840 calculated 
separately for each 
station and unweighted 
average taken 

the monthly mean daily range at Toronto (city) to that at Pearson. These figures were 
rounded to the nearest whole Fahrenheit degree. According to modern practices, the 
"climatological day" for a climatological station that takes only observations in the morning 
and evening ends with the morning observation as far as the daily maximum for the previous 
day is concerned. The daily minimum for Dade observations was calculated as the lowest of: 
(a) the actual evening reading the day before; (b) the morning observation corrected for the 
diurnal minimum; (c) the mid-day actual reading; and (d) the evening actual reading. 
Similarly, the daily maximum for Dade was calculated as the highest of: (a) the morning 
actual reading; (b) the mid-day observation corrected for the diurnal maximum; (c) the 
evening observation corrected for the diurnal maximum; and (d) the morning actual reading 
the following day. In those rare cases where neither mid-day nor evening observations were 
taken, the minimum temperature for the day was calculated as above. Then a maximum was 
estimated using the mean daily range for the month at Toronto (city). This value was checked 
against the morning temperature the following day, and the higher of the two taken. Hence, 
the daily maximum and minimum calculations took under consideration abnormal diurnal 
temperature trends. 

Calculation of mean monthly temperatures. For each month, the mean daily maximum and 
mean daily minimum were calculated from the daily values. The mean monthly temperature 
was then simply the mean of the mean daily maximum and the mean daily minimum. 

Standard difference calculations, Dade minus United States stations. In order to estimate 
Toronto mean temperatures by using Dade data, it was necessary to compare Dade monthly 
means as computed above to those of as many American stations as possible. For each of the 
12 months for the period January 1831 to April 1841, mean differences in monthly 
temperature were calculated between Dade and each of nine American stations: Albany, 
Cortland, Fredonia, Lewiston, New Haven, New York, Oneida, Rochester, and Utica. Data 
for Fairfield, Hamilton and Lowville were not used in this analysis due to many missing 
months of information during the decade. 

Standard difference calculations, Toronto minus United States stations. Similarly, for 
each of the 12 months for the period March 1840 to December 1855, mean differences in 
monthly temperatures were calculated between Toronto and each of the nine American 
stations used in the Dade standard differences above. 

First approximation of Toronto-Dade differences. Since both Toronto and Dade means are 
compared to the same nine American stations, the first approximation of Toronto-Dade 
differences was obtained by subtracting the Toronto standard differences above from the Dade 
standard differences above. These were calculated separately for each of the nine American 
stations, and the overall mean taken for each month. This analysis indicated that, for every 
month of the year, Dade values were high, and that correction values ranging from -0.6° 
(February) to -4.0° (July) had to be applied to Dade means to give a reasonable estimate of 
Toronto means. For the bitterly cold December of 1831, as an example, the Dade calculated 
mean was 15.8°. Since the correction value for December is -2.2°, the first approximation 
of the Toronto mean for the month would be 13.6°F (10.2°C). 

Comparison of mean temperatures for 1831-41 with 1840-55. There was no reason to 

assume that the whole period 1831-55 was climatologically homogeneous. In order to obtain 
a measure of the differences in mean temperature between the early period of the Dade 



observations (1831-41) and the later period of the Toronto observations (1840-55), 
calculations were performed for the three United States stations with the best and most 
continuous observations, Albany, New Haven and New York. Means were calculated for 
each month separately for each of the three stations and for both periods, allowing for those 
months when Dade observations were missing. For each month, an overall mean difference 
(unweighted average of the three stations) between the period means was obtained. The 
earlier period was colder than the later at each of the three stations for each of the 12 
months. The overall monthly differences ranged from 0.4°F (-17.6°C) for April to 2.5°F 
(-16.4°C) for December. 

Second approximation of Toronto-Dade differences. The second approximation considers 
the fact that the earlier 1831-41 period was significantly colder than the later 1840-55 period. 
The first approximation Toronto-Dade difference in mean temperature is -4.0°. The 
difference between the two periods for the same month is 1.2°, so that the total correction 
applied to Dade means to obtain Toronto means is -5.2°F. Because of the variability from 
month to month, Fourier smoothing was applied to the monthly values. As a result, the 
second approximation of Toronto-Dade differences ranged from -2.7° in September and 
October to -4.1° in December. Again, in the case of the frigid December of 1831, the Dade 
mean of 15.8° with a correction of -4.1° results in a Toronto mean of 11.7°F (-11.3°C). 
This is 1.9° lower than the first approximation calculation. 

Preliminary calculation of mean Toronto temperatures using Dade and second 
approximation differences. For each month for which Dade means were available in the 
period January 1831 to April 1841, a Toronto mean was calculated using the second 
approximation differences (Fourier-smoothed) above. 

Comparison of mean temperatures at Toronto by using Dade calculations above and by 
using Method "S". For each month for which Dade means were available in the period 
January 1831 to April 1841, the Toronto mean using the Fourier-smoothed second 
approximation differences with Dade were compared with means as calculated by Method "S" 
(using all available American data). The overall mean differences in the two methods were 
compiled for each month and it was found that Method "S" gave higher values than the Dade 
method in all months - ranging from 0.4° in March to 2.6° in February. The standard 
deviation of the monthly differences between the two methods ranged from 0.6° in June to 
1.7° in January. In the case of the frigid December of 1831, Method "S" indicated a Toronto 
mean of 14.1°F (-9.9°C), 2.4° higher than the preliminary calculation using Dade. 

Final approximation of Toronto-Dade differences. Since Method "S" indicated somewhat 
higher means for Toronto for every month of the year than those by using the preliminary 
Dade calculations, apparently the second approximation allowing for the mean temperature 
differences between the two periods was based on differences that were too great. 
Consequently, the final approximation of Toronto-Dade differences was calculated by 
reducing the second approximation differences by the differences indicated between Method 
"S" and the preliminary Dade calculations above. Thus, the December second 
approximation Toronto-Dade differences, Fourier-smoothed, is -4.1°, the correction due to 
the Method "S" comparison is +1.7°, so the final Toronto-Dade correction for the month 
is -2.4°. Again, because of the month-to-month variation in the correction values, a Fourier 
smoothing was applied. The final approximation of Toronto-Dade differences then ranged 
from -1.1°F in October to -3.2°F in June. In the case of bitterly cold December 1831, the 


smoothed final Toronto-Dade difference is -1.9°, so that when applied to the Dade mean 
of 15.8°, the Toronto estimate works out to be 13.9°, very close to the 14.1°F (-9.9°C) 
in the case of Method "S". 

11. Final calculation of Toronto means. For each month for which Dade means were available 
in the period January 1831 to February 1840, a Toronto mean was calculated using the final 
approximation differences (Fourier-smoothed) above. 

Method "A" 

In the case of Ancaster data, published monthly means based on two observations per day, 9 a.m. 
and 9 p.m., had to be corrected to standard monthly means based on daily maxima and minima. 
Then, standard difference calculations were made between the overlapping records of Ancaster 
and Toronto . From these, estimates were made of Toronto means for those months before 
records began in March 1840. 

Method "A" is outlined in Figure 6. There are three distinct steps: 

1. Correction of mean monthly temperatures. No daily observations are available for 
Ancaster, only published monthly means of 9 a.m. and 9 p.m. observations, from which a 
simple average was computed to produce a monthly mean. Correction values were calculated 
for each month in order to provide monthly means based on the modern practice of using 
daily maxima and minima. These were done by comparing means produced by averaging 9 
a.m. and 9 p.m. monthly means at Toronto’s Pearson International Airport (Atmospheric 
Environment Service 1978) with the monthly means at the same station, which are calculated 
by the usual mean of daily maximum and minimum values. 

2. Standard difference calculations. Ancaster and Toronto data overlap for the period March 
1840 to December 1845. For each of the 12 months during this period, mean differences 
were calculated between the corrected Ancaster monthly means and the official Toronto 

3. Calculation of Toronto means. For each month, January 1835 to February 1840, a mean 
temperature was calculated for Toronto using the corrected Ancaster mean and the standard 
differences above. 

The Reconstructed Toronto Temperature Time-Series 

By using Methods "S", "D" or "A" as appropriate, monthly mean temperatures were estimated 
for Toronto from July 1778 to February 1840. No American data were available for September 
1778 or February, July, August, October, November and December 1779, so that no means 
could be estimated for these months. Beginning with January 1780, a complete set of monthly 
values was obtained. All calculations were done using the Fahrenheit scale, and then the whole 
set was converted to Celsius and combined with Atmospheric Environment Service Archive values 
that begin March 1840 and continue with no breaks until the present day. 

Statistical F tests were applied to monthly and seasonal values to test the homogeneity of variance 

between various periods. In the first instance, three 30-year periods were chosen: (A) 1780-1809; 
(B) 1810-39; and (C) 1840-69. Period A involves only Method "S": through much of this period 








Figure 5: Method "D". 

Original daily data Jan. 1831 - Apr. 1841 
(some months missing) variable hours 
usually 2 or 3 times per day 

Based on mean daily temperature 
cycles for the Toronto area and 
modern "climatological day" practices 

Based on average of daily maxima 
and minima for all complete months 

For each of the 12 months (Jan. 1831 - Apr. 1841) 
mean differences in monthly temperatures between 
Dade (D) and each of 9 U.S. stations 

(Se OCS WS) 

For each of the 12 months (Mar. 1840 - Dec. 1855) 
mean differences in monthly temperatures between 
Toronto (T) and each of the same 9 U.S. stations 
CS) 30% = =e) 

For each of the 12 months mean differences in 

monthly temperatures between Toronto and Dade 

assuming "S'" is the same for both periods 
(CLE = Sy = |p) 


Figure 5: (cont’d) 



1841 WITH 1840-1855 






ap kg 


For each of the 12 months mean differences (C) 
in monthly mean temperatures at 3 U.S. stations 
Period A(1831-1841) from Period B(1840-1855) 

For each of the 12 months difference C 
substracted from first approximation differences 
(Z, = Z - C) (Z, 12-month Fourier smoothed) 

For each month, Jan. 1831 - Apr. 1841, Toronto 
mean (Tp) calculated using Dade mean (D) and 
Fourier smoothed 2nd approximation difference 
(Z,) Gip =O" Zy) 

For each month Jan. 1831 - Apr. 1841 
difference calculated using Dade (Tp) 
fom Method Se (ie) = (Gie— Teas ia) 

For each of the 12 months difference C, 
substracted from Fourier smoothed Z, 
(Z, = Z, - C,) (Zz 12-month Fourier smoothed) 

For each month Jan. 1831 - Feb. 1840 Toronto 
mean (Tp) calculated using Dade mean (D) 

and Fourier smoothed final approximation 
differences (Z,) (Tp = D+ Zz) 








Figure 6: Method "A". 

Published monthly means Jan. 1835 - Dec. 1845 
based on means of two observations per day. 
9° a-m. and! 9) p.m. 

Means based on 9 a.m. and 9 p.m. converted 
to estimated means based on daily maxima 

and daily minima using mean daily temperature 
cycles for the Toronto area 

For each of the 12 months, Mar. 1840 
Dec. 1845, mean differences in monthly 
temperatures between Toronto and Ancaster 

For each month Jan. 1835 - Feb. 1840 


only one, two or three American stations had data available for comparative purposes. Period B 
involves Dade data as well as an increasing number of United States stations. Period C involves 
the early instrumental record at Toronto before urban warming was significant. Only June 
temperature variances are significantly different at the 99% level between periods A and B and 
between A and C. In the second instance, two 40-year periods were chosen: (A) 1801-40 and 
(B) 1841-80. Period A represents the last 40 years of reconstruction prior to the official 
observations beginning in March 1840, whereas Period B contains the first full 40 years of 
instrumental data. Only January variances are significantly different at the 99% level. 


Figure 7: Mean July temperatures at Toronto. 


Individual plotted mean monthly July temperatures from 1780-1870 are shown (Figure 7). The 
cold July of 1816 (the year without a summer) is immediately evident. The trend line is based 
upon a 50-year running mean. 

SE ———,_ 

ae = = 
ane a Fe 
SET Ped stseteenes 
Gis mie 

— Apr 
12 20 
May «— — June 
10 18 
20 22 
July nek — Aug 
Calo i 
16 12 
Peri Sie 
14 | 10 
4 0 
Nov Peles 

Oo i=) 

i=] oS to) [=] o j=] o oO o i=] So oO 
e N & Tt wo © i ioe] o>) i=) = N oO vt wo wo 
© © fo) © © © © © © for) fon) re) for) re) o ror) 
Wo — = 1 f= b = 5 oo r = = — — 7 “_ — 

Figure 8: Trend lines based on 50-year running means of mean monthly and annual temperatures at 
Toronto (values before 1804 and after 1962 are considered constant). 


In Figure 8 trend lines based upon 50-year running means plotted for the middle year are shown 
for all months. Fifty-year means for 1780-1987 can be plotted only from 1805 to 1962. 
Temperatures were considered constant before and after these dates. The overall rise in 
temperature during the period ranges from 0.9°C in January to 3.3°C in October, with an annual 
average of 2.2°C. A significant amount of this increase is no doubt due to the urban heat-island 
effect, which became increasingly significant from the 1880s on. 


Atmospheric Environment Service. 1978. Hourly Data Summaries - No. 3R, Toronto 
International Airport, Ontario. Climatological Services Division, Atmospheric 
Environment Service, Downsview, Ontario. p. 28. 

Craigie, W. 1835. Mean results for each month of eleven years (1835 to 1845, inclusive) of a 
Register of the Thermometer and Barometer, kept at Ancaster, C.W. Clippings from 
newspapers, names and dates unknown. (Unpublished manuscript, Atmospheric 
Environment Service, Downsview, Ontario). 

Dade, Reverend C. 1831-41. Temperatures at Toronto. 47 pp. (Unpublished manuscript, 
Atmospheric Environment Service, Downsview, Ontario). 

Hough, F.B. 1855. Results of a series of Meteorological Observations, Made in Obedience to 
Instructions from the Regents of the University, at Sundry Academies in the State of New 
York, from 1826 to 1850, Inclusive. Weed, Parsons and Company, Albany. 502 pp. 

. 1872. Results of a Series of Meteorological Observations, Made under Instructions from 
the Regents of the University, at Sundry Stations in the State of New York, Second Series, 
From 1850 to 1863, Inclusive. Weed, Parsons and Company, Albany. 406 pp. 

Smithsonian Institution. 1927. World Weather Records. Smithsonian Miscellaneous Collections 
79. The Lord Baltimore Press, Baltimore. 1199 pp. 

United States Weather Bureau. 1932-37. Climatic Summary of the United States; Climatic Data 

Herein from the Establishment of Stations to 1930, Inclusive. Third Edition. United States 
Government Printing Office, Washington, D.C. 


Climate in Canada, 1809-20: Three Approaches to the Hudson’s Bay | 
Company Archives as an Historical Database 

Cynthia Wilson’ 

The Hudson’s Bay Company archives are rich in weather information. Material includes: 
Meteorological Registers; descriptive entries encapsulating the day’s weather or seasonal comment 
in the Post Journals, Correspondence and Annual Reports; and proxy weather data. As a database 
for studying past climate, the strength of the archive is its diversity, permitting cross-checking. 
of results and the convergence of evidence. But the problem of fragmentary evidence has to be 

This paper briefly describes three approaches I have taken to integrate the different material, in 
studying May-October climate during the nineteenth century along the east coast of Hudson/James 
Bay, and over east/central Canada: (1) to establish a detailed regional climatology (1814-21) as 
an historical benchmark; (2) to obtain year-by-year (1800-1900) estimates of monthly temperature 
anomalies (reference 1941-70) and wetness indices; (3) to produce schematic daily weather maps 
(east/central Canada, 1816-18). 

The results of these studies indicate: (1) from 1800-10, May-October temperatures along the east 
coast of the Bay were akin to those today. The seasons then became cooler, with mean 
temperatures falling spectacularly in 1816 and 1817 to values below the modern record; from 
1811-20, they averaged about 1.6°C below the 1941-70 normal; (2) the low temperatures were 
accompanied by greater-than-normal snowfall, and in 1816 and 1817, they essentially precluded 
the growth of many plants. Some areas of snowcover probably remained through the season in 
1816, and offshore, the Bay was barely free of ice at the end of the season; (3) flow patterns over 
east/central Canada from 1 June to mid-July 1816 suggest that spring had been delayed or 
protracted by as much as six weeks; (4) although there was some recovery in 1818, seasonal 
temperatures below the 1941-70 energy level remained characteristic until the 1870s. 

The period of unrelieved record cold from October 1815 to March 1818 may have been 
influenced by Tambora - a relatively short-term volcanic eruption exacerbating a longer-term (60- 
year) lowering of temperature already underway. But there can be no doubt that the exceptionally 
heavy Bay ice lingering through the summers combined with the high frequency of onshore north 
and west winds was a major factor in reducing summer temperatures on the east coast of the Bay. 

Until well into this century, climate loomed large in the daily living and even survival of the 
those inhabiting the Canadian Shield and Prairies, and the Hudson’s Bay Company (HBC) Post 

Journals, Correspondence and Annual Reports provide a rich variety of information directly or 
indirectly pertaining to weather. 

' 90 Holmside, Gillingham, Kent ME7 4BE, U.K. 


The climatic information is of three kinds: (1) Meteorological Registers, often meticulously kept 
in accordance with the accepted practices of the time, but with one or two important exceptions 
on the east side of Hudson Bay, the periods of record are relatively short; (2) descriptive entries 
in the Post Journals encapsulating the day’s weather, with occasional seasonal comment (the latter 
is also found in the Correspondence and Annual Reports); (3) proxy weather information, the 
impact of weather on the natural environment, and on the property, activities and well-being of 
the inhabitants (in all three sources). 

In using this remarkable record as a database to extend the modern climatic series into the past, 
and to study past climatic anomalies, a major problem is that of fragmentation. This has resulted 
from accidents of history, Company policies and activities, and from the nature and interests of 
the individuals recording the events. The strength of the archive is its diversity, permitting cross- 
checking of results and the convergence of evidence. This paper describes briefly three 
approaches that I have used to integrate the material, so as to overcome the fragmentation and 
take full advantage of the diversity in reconstructing (Figure 1): 

1. A regional climatology’, as a climatic benchmark in the historical record. (The east coast of 
Hudson/James Bay, 1814-21). 

2. Extended seasonal time series’ - temperature and wetness indices. (The east coast of 
Hudson/James Bay, 1800-1900). 

3. Schematic daily weather maps’. (East/central Canada, summers 1816-18. This study is still 
in its early stages). 

Some aspects of the reconstructed climate in Canada from 1809 to 1820 have been selected to 
illustrate the rich potential of the HBC archives as an historical database. 

The overall approach to the historical material was traditional, in which the researcher does the 
abstracting so as not to lose vital information offered by the context and subtext. With this in 
mind, weather and proxy data were abstracted in context. Climatically, the historical data were 
approached as far as possible in physical terms, from the standpoint of small-scale climatology, 
the approach of Landsberg (1967) and Geiger (1965). Even with the synoptic mapping, this 
approach was helpful in evaluating and interpreting the individual point data. In this, personal 
experience of several summers in the field at Great Whale’, observing the weather, measuring 
surface energy exchanges and keeping weather journals, has played an integral part. 

Owing to the detailed nature of this kind of work, I do not have space here to substantiate the 
methods or to discuss the assumptions and confidence limits. This information is available, 
together with a full account of the results, in four reports distributed by the Canadian Climate 
Centre (Wilson 1982, 1983a, 1985a, 1988) and three papers published by the National Museum 
of Natural Sciences in Syllogeus (Wilson 1983b, 1985b, 1985c); see also Wilson 1985d. 

! Studies 1 and 2 were carried out under contract to the Canadian Climate Centre, Atmospheric Environment Service, 

Downsview, Ontario. 

The developmental stages of study 3 have been funded by the National Museum of Natural Sciences, Ottawa, as 
part of the Museum’s Climatic Change in Canada Project. 

I am grateful to the Centre d’études nordiques, Université Laval, Québec, the National Research Council of Canada 
and the Canadian Atmospheric Environment Service for the logistical support and research funds which made this 



A. 1814 -1821 


| | | L Jt L = 


PERIOD: 1941-70 


aaa a 
OO S55 








Figure 1: 

B. 1800 TO 1900 


MAY-AUG. 1816, 1817, 1818 





<— BETWEEN ———> 


\ ‘ 



e e 

Studies of past climate in Canada: three approaches to the Hudson’s Bay Company archives as 

an historical database. 

Place and Time of Study 

Considering the amount of work involved, the choice of region and period for study is critical. 
For the regional climatology and the construction of the time series, the east coast of 
Hudson/James Bay (HBC Posts: Eastmain, Fort George, Great Whale, Little Whale River) and 
the active season (May-October) were selected for reasons that follow. 


The Hudson Bay region appears to be particularly sensitive to climatic fluctuations, and the 
eastern windward coast of the Bay provides an excellent laboratory. It is a marginal area with 
respect to the fluctuating arctic/subarctic boundary and the northern limit of tree growth. This 
vast sea, with its seasonal ice cover, open to arctic waters and arctic ice, extends the influence 
of polar climate into the heart of the continent (south of Latitude 52°N) in spring, and remains 
a cold sink in summer; in late autumn the presence of open water creates a snowbelt on this 
windward east coast. To the east lies the plateau of New Quebec/Labrador, a former centre of 
the Laurentian Ice Sheet. All forms of life are so finely tuned to climate along these marginal 
coastlands that any unusually severe or prolonged anomaly can soon disturb the ecological 
balance, and human life and activity - and incidentally make good copy for Journal writers. Other 
reasons for the choice of region include the availability of adequate modern weather and 
environmental records, and my first-hand knowledge of the area. 

The earliest HBC Post Journal for this region is for Eastmain in 1736, but I began the time series 
with the nineteenth century, when better coverage was available for the warm season. With few 
exceptions, there was at least one Post reporting through each season during the century. In 1814, 
the Hudson’s Bay Company gave top priority to a carefully defined program of weather and 
weather-related observations at their Posts in Canada. With wars closing markets in America and 
Europe, fiery competition in the field from the North West Company and a worsening economic 
and social climate at home, the aim was to study and develop the local agricultural potential at 
each Post, to cut the high costs of sending out European food. At a number of Posts, including 
Great Whale, Fort George and Eastmain, fixed-hour temperature and weather data were also 
recorded regularly in Meteorological Registers. The directive fell into abeyance after the 
amalgamation with the North West Company in 1821. Although the Company may never have 
applied its hard-won information, the archives from 1814 to 1821 remain a rich data source for 
detailed regional climatological studies of a period of unusual climatic interest, and were used 

Again, by implementing this programme in 1814, the Hudson’s Bay Company through its 
network of Posts and lines of communication in central, western and northern Canada provided 
a system of synoptic weather observation unique at this time, both in the discipline imposed, the 
consistency of purpose and of manner of observing and recording, and in the extent of its 
coverage. These Company records, together with the logs of the annual supply ships from 
England, coastal shipping and of canoe journeys, offer a basis for synoptic weather mapping. 

From the results of the climatological study for the east coast of the Bay, and given the interest 
in the atmospheric circulation in the years around the eruption of Mount Tambora, the summers 
1816 to 1818 were chosen for initial mapping and investigation. Figure 1 locates all HBC Posts 
with Journals for at least part of the period May-August 1816-18. The density of the network and 
type of weather information available varies through the season and from summer-to-summer, 
depending in part on the regular seasonal operations and needs of the fur trade itself, but to a 
greater extent on the conflict between the Hudson’s Bay and the North West companies. Sadly, 
the battle for the Athabasca trade curtailed weather information from the Red River Valley 
westward from June 1816 through 1817. For the three summers, no alternative historical weather 
sources have been found for the west. For eastern Canada and northeastern United States a 
number of weather records, personal diaries, Mission reports and newspaper articles are available 
for this period to extend coverage (see Figure 1). Regular weather observations were also 
recorded at Godthaab (now Nuuk), Greenland. The search for additional information continues. 


The Three Approaches 

A Regional Climatology, 1814-21 

The prime weather data sources in the HBC archives are the Meteorological Registers, even 
where they were kept for only a few years. To try to make full use of them, one approach is to 
analyze and integrate all aspects of the local or regional weather from all HBC sources for those 
years - temperature, winds, cloud, precipitation, extreme events and, rarely, pressure, together 
with all proxy-weather indicators - to gain as clear a picture as possible of the climate of that time 
in terms of the present-day climate: that is, to set up a climatic benchmark. 

The dangers of this approach are only too well-known: the differences between the historical and 

modern instrumentation, exposure, observing practices and so on. Happily, a tradition of careful 
meteorological observation and reporting had become established on the west side of the Bay in 
the second half of the eighteenth century with the collaboration of the Royal Society, which 

advised on instruments and procedures. This tradition continued into the early nineteenth century. 

Provided that basic assumptions are made explicit and their import is clearly stated, and that 

every effort is made to compare like with like, I believe the results to be worth the time and 

effort required. 

With the historical temperature readings at Great Whale/Fort George and Eastmain, I tackled the 
calibration from several directions, hoping in this way to approach a consensus and to avoid 
circular arguments. The three main lines of attack were: 

1. Historical - the reconstruction of the early observing sites, and social context, and of the 
meteorological instrumentation and procedures accepted at the time. A study was also made 
of the history and. homogeneity of the respective modern temperature records. 

2. Physical - examining the systematic temperature differences that might arise from changes 
in site, instruments and their exposure, and observing practices, given the distinctive qualities 
of the subarctic surface conditions and regional and local weather. 

3. Statistical - an application and extension of current Canadian quality-control procedures, the 
fitting of simple regression models, and the analysis of the fields of error. 

Although corrections were made to the daily maximum and noon temperatures, and also to the 
daily minimum where the start of the climatological day differed, I was impressed by the 
consistency of the historical temperature record in the context of the instruments and procedures 
of the time. 

Extended Seasonal Time Series, 1800-1900 

The basic HBC sources of continuous and consistent weather information over extended periods 
are the descriptive entries in the Post Journals. A reading of the Journals often leaves a strong 
impression as to the relative heat or cold, dryness or wetness of the different seasons, through 
the subtle integration of the many different weather and environmental factors and their impact. 
Thus a second approach is to try to integrate daily and seasonal weather remarks and all forms 
of proxy-weather information, to obtain monthly estimates of the temperature anomaly with 
respect to a modern reference period (in this case 1941-70); also, to obtain monthly wetness 
indices with respect to the modern precipitation record. The benchmark set up in the first study 
acts as a uSeful reference for the early part of the century. 


The approach is similar to that taken by Pfister (1980) in Switzerland. For the thermal series, a 
first approximation to the monthly temperature anomaly is obtained from the direct weather 
remarks, then the proxy indicators are applied to try to obtain an order of magnitude (timing, 
intensity, duration), or at least supporting or modifying evidence. One major difference here is 
that a greater variety of information must be used to compensate for the fragmentation of 
individual data series, to give a convergence of evidence. 

As the method must accommodate so many different kinds, fragments and combinations of data, 
some firm climatological structure is required. A secure, yet flexible modern frame of reference 
was provided by the nine modern daily temperature curves for Great Whale, Fort George and 
Eastmain, respectively - comprising for each day of the May-October season the reference period 
daily mean, the highest and lowest daily mean, the daily mean maximum and the highest and 
lowest maximum, the daily mean minimum and the highest and lowest minimum. Other aids 
included a wide variety of modern temperature and temperature-related information, including 
analogues for warm and cold months. 

The proxy-weather sources (Figure 2) can be grouped into three: snow and ice, phenological 
events (plants and animals), and human activities. These data have been approached from two 
points of view: 


1 Ons 
a 5G NG A 
—---- OG OCcC--—----- 



' WTH 

Figure 2: Summary of environmental indicators. 


1. To obtain an indication of the timing and magnitude of seasonal and unseasonable events, and 
to compare where they intersect the modern daily or weekly temperature curves. One avenue 
of attack is the heat-unit concept, with thresholds 0°C, 5°C and 10°C; another, the cardinal 
points with respect to the different crops; a third, the agroclimatic capability classes, which 
delimit in climatic terms the crop potential of the region - and so on. 

2. To try for a statistical link between the proxy data and the monthly temperature anomaly 
during the modern period of record, to provide guidelines or rules of thumb. 

All the information was integrated for each season and each Post in direct comparison with the 

modern reference period to give the temperature anomaly month-by-month from May to October. 

Where more than one Post was reporting, the anomalies were then compared and combined. The ' 
period of greatest confidence is 1815-20. 

The wetness index is a five-point scale based on the number of days with reported precipitation, 
together with supporting remarks and proxy evidence indicating its intensity or duration, and the 
degree and duration of dry periods. Each month and each season, for each Post, was assessed 
directly against the modern record of precipitation at the respective weather stations. Following 
Pfister, the upper and lower quartiles were chosen as limits between wet and dry months; also, 
the octiles define very wet and very dry classes. Adjustments were made to account for those 
days when rain fell only at night and may not have been reported, and to account for discrepancy 
in the modern reporting of the number of days with snowfall between 24-hourly observing 
stations and climatological stations (cf. Ashmore 1952; Manley 1978). 

Schematic Daily Weather Maps, Summers 1816-18 

A third approach permitting the integration of all available kinds and fragments of weather 
information is synoptic mapping, although the Meteorological Registers with their detailed timed 
weather observations act as linch pins in the historical analysis. 

Since the historical data include few records of atmospheric pressure at this time, these weather 
maps for east/central Canada are primarily based on surface wind data - more readily and 
frequently recorded. Thus they offer schematic representations of the flow patterns, rather than 
the refinement of Kington’s (1988; this volume) classic series of daily weather maps for western 
Europe and the northeastern Atlantic in the 1780s, and in 1816, which are firmly based on a 
network of barometer readings. That surface winds can be so used has already been elegantly 
demonstrated by Lamb (in association with Douglas) for the period of the Spanish Armada, May- 
October 1588 (cf. Lamb 1988). Caution is required. Regional wind direction and speed can be 
modified at some sites by local topography, the geometry of forest and other obstructions, or 
masked by the influence of local sea or lake breezes and valley winds. But marked local effects 
can usually be detected and allowed for if the network is reasonably dense and given some 
knowledge of the terrain. 

As a first step in the map analysis, all the direct and proxy-weather information are plotted on 
daily base maps. Using transparent overlays, the data for morning and evening hours are 
transposed to separate charts. Each chart is then analyzed over the base map, which provides the 
necessary background information as well as intermediate history, and in conjunction with 
previous maps. Areas of cloud and precipitation are shaded in, the temperature and wind fields 
studied, the zones of maximum gradient, wind shear and the pressure tendencies noted. Frontal 
zones are tentatively indicated. Then an attempt is made to sketch in the pressure pattern, bearing 
in mind the wind speed, the nature of the surface and the most likely direction and speed of 


movement of the frontal systems. As an aid to analysis, where Registers exist, daily temperature 
curves have been drawn up by month, with winds, cloud and precipitation added, to give a visual 
display of the sequence of weather. All available pressure traces have also been plotted. 

Here, the question of data calibration is partly resolved by the space and time smoothing implicit 
in the map analysis. With respect to temperature, it is the relative differences and changes that 
are important, and any significant errors would be expected to stand out. The results of the earlier 
calibration study of the HBC data suggested that observing practices at this time were consistent 
throughout the network; also that temperature readings were most reliable in the morning and 
evening and near freezing. The wind-force scale in use on the Bay was quite similar to the 
Beaufort Scale, and the latter has been used to convert all indicators to approximate speeds. Wind 
direction was generally assumed to be with respect to true north. The barometer readings have 
been reduced where possible to sea level. 

Concerning the synchronization of data across Canada, the Hudson Bay region is considered as 
the reference time zone; the western and eastern extremities of the map are then within about + 
two hours, which can be born in mind. The timing and frequency of the observations are seminal 
to the analysis. Problems here can often be overcome when maps are sketched in for both 
morning and evening. In drawing up early weather maps, “historical continuity" becomes a prime 
tool in grappling with the many difficulties, including that of sparse data. There is also the 
advantage of knowing (if only in part) the future as well as the past. 

Notes on Climate in Canada, 1809-20 

What can be gleaned from these studies about climate in Canada from 1809 to 1820? Are there 
any clues following the major volcanic eruption of Mount Tambora in April 1815 as to the 
possible influence of such massive atmospheric loading on regional climate? To what extent was 
the atmospheric circulation over east/central Canada anomalous at this time? The following notes 
illustrate some of the kinds of climatic information that can be reconstructed using the Hudson’s 
Bay Company archives as an historical database. 


To set the climatic events of the period 1809-20 in context for the east coast of Hudson/James 
Bay, Figure 3 shows the reconstructed series of May-October mean temperature anomalies 
through the nineteenth century together with the modern twentieth century record. The reference 
period is 1941-70. During the first decade of the last century, mean temperatures were similar 
to those today, but then the seasons quite rapidly became cooler, tumbling in a spectacular fall 
in 1816 and 1817 to values far below the modern records. Seasons well below the 1941-70 
energy level remained characteristic of the first half-century with a cluster of cold years from 
1835 to 1840; decadal averages for 1811-20 and 1831-40 were about 1.5°C below. A certain 
mid-century amelioration was followed by some very cold weather in the 1860s, before a 
remarkably sudden change occurred in the early 1870s to a new higher energy level akin both 
to that at the beginning of the nineteenth century and to that today. Although not as warm as the 
1870s, the last two decades sustained these milder conditions, and variability remained well 
within that for 1941-70. The upward trend from 1811-40 through 1870-1900 shows clearly in the 
overlapping 30-year means. This sequence of events has no analogue in this century. The volcanic 
eruptions of Tambora and Coseguina were followed by extreme cold, but the decline had set in 
beforehand. For Krakatau and Agung, any similar signal (if that is what it is) is weaker, and for 
St. Helens and El Chichén, absent. 


(+0.11) (-0.10) (-0.07) (-0.08) (—0.03) (-0.04) (-0.06) (+0.13) (-0.01) (+0.07) 



Figure 3: East coast, Hudson/James Bay. Reconstructed mean temperature anomalies (°C) summer 
seasons (May-October) 1800 to 1900. M, decadal means; superscripts, number of years. Tr, 
maximum tree-ring density, decadal mean anomaly (Parker et al. 1981). Anomalies with 
reference to the 1941-70 period. 

Within the May-October season, the most striking feature of the nineteenth century is the coldness 
of spring (May/June) and autumn (September/October), and the effective shortening of the active 
season. The curve for the spring months closely parallels that for the season as a whole, but in 
autumn, cold seasons persist through the last two decades, although not below the modern record 
low of 1974. In complete contrast, midsummers (July/August) were in general not so very 
different from those today, although 1816 and 1817 with anomalies of about -5.5°C and -4.0°C 
were in a class by themselves; cold summers also occurred in 1836 (-2.5°C) and 1871 (-3.0°C). 
By far the coldest summer on modern record was 1965 (-3.5°C), two years after Agung. 

Focusing on the second decade of last century, with its remarkably sudden lowering of 
temperature, Table 1 gives the absolute mean temperatures for 1808-20 (estimated for Great 
Whale from the reconstructed regional anomalies) together with the 1941-70 normal values. The 
first strike in the change to a colder mode was winter 1808-09, heralded by a very cold autumn 
in 1808. Gladman, Master at Eastmain (a keen observer and a veteran of these shores), found it 
the coldest winter he had ever experienced, with temperatures frequently below -40°C and 
scarcely any mild weather. The next event was the extremely early and very cold autumn of 
1811, advent of the even colder longer winter of 1811-12; the two spring months May/June 1812 
probably averaged below freezing at Great Whale (Table 1). These were considered extraordinary 


M (+0.3) (-16) (-1.1)8 (-1.5)§ (-0.8) (-0.6)° (-0.9)° (+0.1)? (-0.7)8 (-0.2) 
1800 1810 1820 1830 1840 1850 1860 1870 1880 1890 1900 
SC Le ee ee ee ee a a | a | eee | a ee ees | eee een (oy 
+3 +3 
es +2 
+1 J [] | 2 +1 
4 } | | f i | 
oF Oe ren) QT ro) i: 
al U U1 | 1 | | | =! 
U ]] | | 
-24 | | l -2 
-3- 1} | -3 
-44 U | -~4 
a5 | -5 
-~6 -6 
1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 i 
be] | ° 
+35 +3 
+244 +2 
+1 i] +1 
0+) LT. fo) 
=j(=' -1 

times. In 1811, the HBC annual supply ship from England did not arrive at Moose Factory until 
25 September, which was so late in the season and unprecedented at that time that there had been 
great alarm at Eastmain lest no ship should arrive. As Gladman wrote: "these circumstances are 
altogether so new and unfortunate" (HBC.B59/b/30). When the ship sailed for England, it could 
no longer leave the Bay for ice, and wintered in Strutton Sound off Eastmain. An HBC supply 
ship had last wintered over (though storm damage) in 1715 (Cooke and Holland 1978). It was 
to happen again, as a result of ice, in 1815, 1816, 1817, 1819. 

The most persistent and intensely cold period, continuously below the reference normal, began 
in autumn 1815 and lasted until late winter 1818. The degree of cold reached its nadir in 
January/February 1818, to be followed by a remarkable flip to an early, warm spring and a 
benign season. Following this, cold springs and autumns continued. For 1815-20, Figure 4 shows 
the monthly temperature anomalies for May-October at Great Whale, Fort George and Eastmain - 
adjusted values based on temperature observations in the Meteorological Registers (cf. Wilson 
1983b). Looking more closely at these seasons in Table 1 and Figure 4, several features are 

1. The degree of the anomalous cold in 1816 and 1817, with many of the months below the 
modern record - Alder, Master at Great Whale, then at Fort George: "if summers I may call 
them" (HBC.B77/e/1a). At Great Whale in 1816, July appears to have been nearly 6°C below 
normal, that is 2°C below the lowest on record 1965. The modern standard deviation is 
1.2°C. At Fort George in 1817, the season as a whole was some 5°C below normal. As a 
season 1817 was more severe than 1816, but the greater severity was in spring and autumn 
rather than midsummer. 

2. From autumn 1815 through April 1818, the greater part of this coast experienced arctic 
conditions following K6ppen’s definition. During summers 1816 and 1817, the arctic/subarctic 
boundary lay close to Eastmain, some 3.5° latitude south of the present average position near 
Richmond Gulf. The closest modern analogue is probably the 1965 season. 

3. In 1815 and 1816, the rise of the daily mean temperature through 0°C, the start of the active 
season, was two to three weeks later than the 1941-70 normal, akin to 1972. In 1817, it was 
some four to five weeks late and the return through 0°C three weeks early in autumn, hence 
the period above 0°C was some two months shorter. This was reflected in the break-up and 
freeze-up of river ice. It is almost certain that seasonal ice remained in the ground in many 
areas throughout the 1816 and 1817 seasons. It was also reflected in the snowcover, and type 
of precipitation. The implications with respect to plant growth can be clearly seen in Table 1. 


Again, to set the 1809-20 period in context, Figure 5 shows the series of May-to-October wetness 
indices for the east coast of Hudson/James Bay through the nineteenth century, and the modern 
record expressed in the same form. In general, the first half of last century was not only colder 
but wetter than today, while the second half became warmer and drier. By far the wettest decade 
was 1811-20, with a run of wet seasons from 1814 to 1820. The three wettest seasons of the 
century 1816, 1817, 1820, at least matched the record in 1944. The three driest seasons were 
1809 (which probably equalled the record 1920 season), 1807 and 1878. The rapid change from 
the warmer/drier mode of the first decade to the very cold/wet regime of the second is 


‘poul[iopun oie spotied pjoo Ayjensnuy, , 



0} Ae 

€lsl | CIsl 6081 | 8081 

07-8081 (DO, Ul senyea aynjosqe) seinjesodule], URa] payonsjsuocey ‘yey AH :] 3quL 



g LO[MIYidAISIO[M| Jj JjajsjolMj J) Japs jo[MjJ{JpAjsjolMjyjJjAsjolMjJiJpaisjolmis| 
|_ WR BR | 
= 1815 1816 1817 | 1818 1819 1820 1821 
2 r2 =e 
6 ™ MM (a MM i [_hpegister missing | > 
im ai lh Taal = TI ae =| Saal 
- | ie | =| 
c° c 
2b | = 4-2 
4 -4 
6 ‘iz -|-6 
— below modern = 
2) - all = 
= | FH 
0 ss) + + T | SSS => ae ft malteaenl| © 
ale Sana 2) IL | | | 4 
“A je | a * ioe | 2 
ei a 
i= | L* | * | | | Se! 
lEW |e ee } 4-4 
Come ie ae = |) GRC: 
=6\ |= * + -6 
L naif 1 
alt 1815 1816 lex tsi 1818 1819 1820 1821_|_ 
= * | 

* Below the short 1960-72 record 

Figure 4: Whale River (Great Whale), Big River (Fort George), Eastmain, 1814-21: mean daily 
temperature (adjusted values) expressed as differences from the 1941-70 normals (WR, BR) or 
1960-72 averages (EM). The shading and asterisks indicate where the historical mean was 
below the extreme monthly mean on modern record. M, data missing. These monthly 
anomalies, together with the absolute values, are tabulated in the Appendix. 

(-3) (+9) (+2)? (+1)? (+3) (-3)® (-3)° (iy (-3)° (-7) 
1800 1810 1820 1830 1840 1850 1860 1870 1880 1890 1900 

fal’ 2 


io) (e) E= 
S Oe 

1900 1920 1930 1940 1950 1960 1970 1980 1990 2000 

+4)’ o (+1)7 
3 1 
Vw Sal |e VW 
Ww Ww 
fo} 0 
: | | ! 
1 3 


Figure 5: East coast, Hudson/James Bay. Wetness indices, summer seasons (May-October) 1800 to 1900. 
Index from +2, very wet to -2, very dry. In parenthesis, decadal sums of the index; 
superscripts, number of years. Asterisks indicate borderline cases, wet or dry; 1,2,3 wettest or 
driest season. Reference periods: Great Whale, 1926-76; Fort George, 1916-69; Eastmain, 


Looking at the detail in the Meteorological Registers and Post Journals for the seasons 1815-20, 
the effect of such low temperatures on precipitation is evident. Figure 6 illustrates the greater 
number of days with snowfall from May through October, and the shortening of the snow-free 
season, contrasted with the modern period. The effect is especially noticeable on James Bay, 
suggesting southward extension of the autumn snowbelt. Of particular interest is the summer 
snowfall in 1816 at Great Whale; not only was there more snow in July than today, but even 
more fell in August - a month that has no modern record of snow having fallen. Summer 1816 
provides a marginal case for a residual snowcover on the east side of the Bay. In summer 1817, 
snow conditions at Great Whale were most likely even more extreme. At Fort George, snowfall 
was extraordinarily frequent and often heavy in May and June 1817, but no snow fell in July and 
August, and the heavy rains of August probably washed away any snow remaining at the coast. 

Thus the May-October seasons of 1816 and 1817 were such that had these conditions persisted, 
they might have resulted in the formation of permanent snowfields in parts of New 
Quebec/Labrador. From the impact of these seasons recorded in the Post Journals, 
Correspondence and Annual Reports, it was indeed possible to see the southward expansion of 
snow and ice forcing back the northern margins of habitation along this coast. The association 
of volcanic activity and incipient glaciation is an old idea in the literature of climatic change. 

GREAT WHALE (1941-70) FORT GEORGE (1941-70) EASTMAIN (1951-80) 
12 12 

2 11) MAXIMUM 




BR 16 
” 1817 1818 1819 1820 1821 | » 
6 12 6 
: alas 
i 4a G 
: : 
: | sNeree 
oe J : MM | J JJ] spud J FF 9 
|Slo[MlyTuTATsto[Mlulstatstol[mlulstatsiol[mlslstaTslo[mMlulslalsto[ml J STatstoimtata| 


ye [1814 1 é 
i =e 
ie 1815 1816 1817 1818 1819 1820 1821 o 
< Alvi ie 
ie 412 0 
6 4 6 
& BE 4s G 
o | a 
= +b Slee 
2) = 
z 4e meh Fe 


0M |w_ wf} MM | J 

slo[mlulJTAlSlo|mlslslaTSlo[mlalatalsto[mlulglals o[w JIJTATSTo[ml ulutalslofmlsls] ° 

Figure 6: Whale River (Great Whale), Big River (Fort George), Eastmain, 1814-21: number of days with 
snowfall, together with modern reference values for Great Whale, Fort George and Eastmain. 
(The climatological day beginning at 8 a.m. at Whale River and Big River, 6 a.m. at Eastmain.) 
M, data missing; J, Journal entries, no Register; *, less than 1 day. 


Regional Climate and Tambora 

Circumstantial evidence for the east coast of Hudson/James Bay from autumn 1815 to later winter 
1818 suggests a possible case for regional climatic cooling through the intervention of Tambora. 
That the material from the equatorial eruption in April 1815 should have entered the polar 
stratosphere by autumn of that year, with a residence time of more than one year, is in keeping 
both with the structure and behaviour of the atmospheric circulation and with studies of 
radioactive fallout in the 1950s and early 1960s. Moreover, empirical studies and certain 
theoretical considerations suggest that any resultant lowering of air temperature near the surface 
in higher latitudes might be expected to be most apparent in the warm season, and of greater 
magnitude than in lower latitudes. But in the case of Tambora (and of Coseguina), this appears 
to be at most a short-term feature superimposed on longer-scale climatic changes. The onset of 
cooling in this region occurred before the major event of Tambora (cf. Figure 3), with the return 
to a warmer mode some 50 years later. While a number of smaller eruptions did take place in 
the years preceding Tambora, there is also the concurrent event of the double cycle of abnormally 
low sunspot number (the lowest since the Maunder Minimum), which spanned the first two 
decades of the nineteenth century (Eddy 1976, p. 1191; this volume). A further consideration is 
the sudden swing to near-record warmth in April/May 1818 following hard upon the coldest 
weather recorded; this almost suggests an "over-compensation" in redressing the balance. Had 
there been a significant scavenging of the aerosol by the end of the very wet 1817 season? 

For a different perspective on the 1816-17 seasons, it is useful to consider the regional climatic 
controls and energy exchanges - in so far as clues are offered in the HBC archives - without 
directly invoking Tambora. 

Energy Exchanges: Advection 

The thermostatic effect of the Bay on summer temperatures is a critical, if complex, influence 
along this windward coast, where background air temperature level appears to be closely related 
to the temperature of the Bay surface. Cold seasons do tend to have a higher proportion of Bay 
winds, and this was the case in 1816 and 1817. 

The arctic summers of 1816 and 1817 were marked by heavy ice and late break-up and melt in 
the eastern and southern parts of the Bay. In both years, the last heavy ice was compacted in 
southeastern Hudson Bay, extending into northern James Bay - a pattern similar to modern 
maximum ice/water limits for mid-August to mid-September. In 1817, the timing was perhaps 
a week later. But in 1816, these stages occurred some four weeks later between mid-September 
and mid-October. In mid-September 1816, the ice in James Bay was akin to the normal for mid- 
July. Given the cold autumn of 1816, the season provides a marginal case for the carry-over of 
ice from one season to the next - a situation exceptional for the period itself. In contrast, the 
clearing of the ice for the 1818-20 seasons was relatively early in this part of the Bay. 

The windroses for Great Whale, June to August 1816, are shown in Figure 7, together with those 
from the modern record. North winds are normally frequent here in May and early June, 
associated with a series of anticyclones which cross from the Arctic into southeastern or eastern 
Canada; this is then superseded by prevailing upper westerly flow with travelling depressions, 
which is characteristic of summer. Figure 7 shows how the spring pattern continues through July 
in 1816 with an unusually high frequency of north and west winds. In August the pattern 
changes, but winds are overwhelmingly from the west. These months are dominated, then, by 
the advection of cold air either from the Arctic or from passage over the Bay ice, and by the 



1816 1942-54 
N (RIVER BANK sites 1, 2) 
30% NO% 1967-76 
JUNE 20% JUNE 20% 
10% 10% 
30% 20% 20% 10% 10% 
ne rae Calm 3% Calm 6% 
== 30% 
40% 30% 20% 20% 10% 20% 10% 

No calms Calm 2% — Calm 6% 
N N 
50% 40% 30% 20% 20% 10% 20% 10% 
Calm 2% Calm 5% 

No calms 

Figure 7: Windroses for Whale River (Great Whale) 1816, and Great Whale 1942-54, 1967-76. (The 
recent river bank sites are comparable with that of 1816.) 


> > > 


5 3 3 

@ MW rf 

ea a eq 

w uw w 

kb ie = 

Zz Zz Zz 

WwW w Ww 

12) Oo 12) 

c fog joa 

Ww Ww Ww 

a a a 

}<—WwR—>| BRt+> 
3B 60) = 1815 1816 | 1818 }— 1819 1820 18214 60% 
:) 4 
ae ey 4 4 
z le} 
3 40 + 4 40 0 
re se 
= z 
% 20+ + 20 & 
iva tL Sh Shy 
i No data M MM No Register a 
e TuTyTal TalslolMlulal Talslolmlu| ° 
O[mTaTyTaTsTo[Mlulutatsto} JTATSTO[Ml JT STalsto[ml sTuTatsto[ml 

1815 1816 1817 1818 1819 


Ic |M MMM MM | 


Figure 8: Whale River (Great Whale), Big River (Fort George), Eastmain, 1814-21: relative frequency 
of "clear" hours, together with reference values for average, warm and cold months at Great 
Whale. Morning, noon and evening hours. (A “clear" hour is defined by zero to five-tenths 
cloud cover.) 


almost total absence of the warmer southerly or land components. In 1816, as today, snow in 
spring and early summer was brought by northerly and westerly winds. The persistence of 
onshore winds in turn served to pack the ice in along this coast all summer, and further depress 
coastal temperatures. At Fort George and Eastmain in 1817, the prevailing onshore winds from 
June through August indicated the frequent passage of depressions. This suggests that the 
exceptional cold of these two seasons was associated with different circulation systems. 

Radiative Energy Exchanges 

Today, low cloud is dominant in this region during the average summer season, and in 
exceptionally cold summer months even more pronounced (Figure 8). In sharp contrast, a striking 
feature of the seasons 1815-20 in general is the greater frequency of clear weather (zero to five- 
tenths cloud cover) from May to August. In 1816 at Great Whale, the frequency was double what 
might be expected today in very cold months, and particularly noticeable in July. Respective 
listings of clear hours against wind direction and damp, cloudy hours with Bay winds suggest, 
when compared with modern analogues, that the clear weather in 1816 was the result of: (1) the 
prolonged influence of arctic airmasses at this period; (2) the greater frequency, persistence and 
intensity of spring anticyclones over Hudson Bay, probably extending through July; and (3) the 
late break-up and unusual persistence of heavy ice in the Bay through the summer. 

Still leaving aside Tambora, the clearer skies in 1816 and 1817 compared with very cold summer 
months today imply a larger receipt of incoming solar radiation at the surface, although the full 
potential may have been reduced, particularly in the region, as a consequence of the unusually 
"quiet" sun’. Considering the short-wave radiation balance, any increase in incoming radiation 
at Great Whale in 1816 could have been more than offset by increased losses resulting from the 
exceptional clarity of the air and through reflection from late snowcover and ice, enhanced into 
July and from the third week in August by fresh snowfall. In the case of the long-wave radiation 
balance, the dryness and clarity of the air and low sky temperatures would have encouraged loss 
from any more favoured sites or surfaces, while the net radiation through the summer would have 
been used primarily in melting snow and ice, and thawing and drying out the soil. These 
conditions together with the low-level advection of cold air could go some way to account for the 
very low air temperature at screen level at Great Whale in summer 1816. 

Reintroducing Tambora, measurements of solar radiation following the eruption of El Chichén 
in 1982 suggested a reduction in the total short-wave radiation reaching the surface as a result 
of the stratospheric loading of dust and sulphur; while there is satellite evidence of enhanced 
infrared emission from the cloud, the effect of the volcanic material on the long-wave balance 
at the surface is not known. 

To speculate, the evidence so far suggests that the cause of the extreme cold along this coast at 
this time was most likely multiple: a combination of unusual external factors converging on the 
years 1816 and 1817 (the general decrease in solar power, coupled perhaps with a high frequency 
of heavy volcanic aerosol in this particular region of the stratosphere), whose climatic effects in 
"summer" were magnified along this subarctic/arctic margin through the massive presence of 

' Given the apparent connection between auroral/geomagnetic activity and sunspots, it is worth noting that compared 
with the high frequency of auroral activity observed in recent years, the only reports of auroral sightings in the HBC 
Journals for this coast during the nineteenth century were in 1878, 1879 and 1880 (cf. Figure 3). 


unusually late ice’ and snow, and the complexity of the ensuing surface - atmosphere 

Atmospheric Circulation, June-July 1816 - Preliminary Remarks 

The maps from the sequence 1 June to 13 July 1816 are first approximations to illustrate the work 
in progress (e.g., Figure 9). The results of the pilot study for 1-17 June (Wilson 1985Sc, 1985d) 
had indicated that useful daily schematic flow patterns for east/central Canada can be drawn for 
this early period, with the HBC archives providing the core database, supplemented where 
possible by other historical weather sources. Two sample maps are reproduced here (Figures 9a, 
b); additional information obtained more recently for the east coast of the United States and 
Canada (cf. Figure 1) is now serving both as a check on the original analysis and to refine the © 
patterns. Although the HBC data are less complete in July, the coverage is still adequate when 
the weather patterns are well-articulated, which was generally the case in the first two weeks 
studied to date. At this early stage of the study, one or two preliminary remarks can be made 
concerning the atmospheric circulation during the first half of summer 1816. 

During much of this period, flow over east/central Canada and the northeastern United States was 
predominantly meridional, interspersed by short periods of more zonal flow (notably the first 
week in July) with rapidly moving depressions, and brief northward extensions of the Subtropical 
High. The synoptic situation for the period 5-10 June points to blocking in the vicinity of Hudson 
Bay; from 6 July until the end of the present analysis on 13 July, there is some evidence, 
provided by the approaching HBC ships, of a blocking high east of Greenland. 

The two exceptionally cold events over eastern Canada and the United States (5-10 June, 
6-11 July) were apparently associated with these periods of blocking. In each case, a depression 
passed across the Great Lakes/northern Ontario (cf. Figures 9a, 10), lost speed abruptly over 
Québec and developed into a large system, gradually drifting eastward. Behind the depression, 
high pressure extended from the Arctic down over Hudson Bay, and very cold air was pulled 
unusually far south in the rear of the storm. In June, the bitter northwest winds brought frost and 
snow to the St. Lawrence Valley and New England (Baron, this volume). In July, the clear dry 
air brought very low temperatures, especially at night, at least as far south as Philadelphia, where 
it was as cool as late September - and mornings and evenings uncomfortably so’. 

The intensity and size of some of the systems, as well as the highly variable and contrasting 
extremes of temperature experienced throughout the region, bear witness to the vigour of the 
north-south energy exchanges, and suggest a much stronger mid-latitude temperature gradient 
than is usual today at this season. The storm tracks in early July were unseasonably far south. 
A strong temperature gradient was present at times to the south of Hudson/James Bay between 
the forested/spaghnum Shield country north of the Great Lakes, which on occasion became 
extremely warm, and the unusually complete and compacted ice-covered surface of the Bay; here, 
lows seemed to regenerate or develop. 

' The possibility of submarine seismic activity in the Arctic as a source of kinetic and heat energy, easing the breaking 
up and outflow of previously compacted arctic ice, has still to be ruled out (cf. Wilson and MacFarlane 1986). 

2 Deborah Norris Logan’s diary, Historical Society of Pennsylvania. 


998 34) 
rising’ snow 

ne Bor overcast 

(No geese flying) ~ 

. . 
very fine morning 

¢ 3 . severe frost 
\ x snow last night | 

: | very cold, snowing) f W Gey ss 
~ : frostl. 

ste See il tenement” stougy, \F | 1¢ [iMfenowing 7738) 
(killed plants, singed oak leaves) ¢ ground)¢ snow! He ial eee light rain 
Seep Ae) very hard frost os 

\colds* Gat 
arg Neo cloudy Dies 

*~ cenOw, on around) 
: N\ 32 
. snowing \(33) : 
Ncloudy 46 
S ae 

SLE 60 cold, unpleasant day 
June 5 1816 and sultry) vineavy rain (rain began 2p.m.) 

———~. Schematic pressure patterns OS fain 
Fronts: <A, warm, ~~ cold 

° 200 400 800 800 1000 Lenin | 

— a nd kL OMETRES 

100W 70w 

Figure 9a: Surface weather map, 5 June 1816, morning. Temperatures in degrees Fahrenheit; in 
parenthesis, mid-day values. Winds, short barb five knots, long barb 10 knots; broken arrow, 
one observation a day, time unknown; asterisk, speed unknown. Pressure in millibars; recent 
work has suggested that an adjustment of about +9 mb is required to reduce the station 
pressure at Québec City to sea level. (Reproduced with kind permission from Weather 10, 
Pp. 1375) 


18 995 
(57) \ falling 
calm with rain 



= SS 

f (60)7~ overcast, 
heavy rain 

(Bad weather @ 

prevented embarking) —; 33 Train at 

(57) times\: 

calm clear 

rain, thunder andle (55) err Gan day) 
Pon tningn@ S 

6) 91011" % 
Coes steady, \ 

June 10 1816 33 ’ 
Mornin (60) 7 Clear 
me o (water froze) 
(60) 1018 
clear@rising | 
° 200 400 600 600 1000 

pe ot rl 1 OME TRES. 


Figure 9b: Surface weather map, 10 June 1816, morning. For legend, see Figure 9a. (Reproduced with 
kind permission from Weather 10, p. 137.) 


= ay =O is : ~ 4 , 
7 es, oe \ 

: June 10e | \ 
June 10 m @ Ne \ 
; ® © } \ 

June 9 m\ | \ 

| ‘\ \ 

| \ \ 

\ \ 
} x | 
\June 9 e @June 9 e 
~ oe June 13 m », 
ao : : Ps : Na @ ® June 9 
: é wo June 11 e® ® June 12 e ~ 
‘ S : Gs a June 12m}. June 8 e \ 
~ : = a oat \ 
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~ ‘ : \ a \ June 7 &—June 8 m \ 
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= : : DX NK June 6 e/“June 7 m 
om : \ June 5e —@—~ 
eee N June 6 m 

June 3m: r 4 ‘ \ oe 
= as june 4m : @June 5 m ———= 
ay x June 4 e “oJune 10m 
“. ; \ 
ew ial : 
: June 10 e' wy 
oe i. 4 fy W—7 ["s 

June 3 to 13, 1816 
@—— Lows \ \ 

@ ~~ > Highs fas \ ay 
m. morning, e. evening \ - 

° 200 400 600 800 1000 L— A . — \ 
———— — KILOMETRES a June lie 

100W sow BOW TOW cow 

Figure 10: Trajectories of the surface high- and low-pressure systems, 3-13 June 1816. (Reproduced with 
kind permission from Weather 10, p. 138.) 

It appears that the stalled situation of 5-9 June was associated with the intensification of an 
anticyclone over the Bay ice. On 9 June, the high began to move out to the southeast (cf. Figures 
9b and 10) and was centred off the coast of New England by the evening of 11 June. Although 
less persistent, this anticyclonic pattern of flow was repeated in the third week of June - the 
trajectory now more southerly, over Lake Erie. The analysis is not yet complete, but a rather 
similar situation may also have occurred at the time of the stalled circulation in July. Today, the 
Hudson Bay High is characteristic of spring, (March-May) when anticyclones are more frequent 
over eastern Canada. Although some of these spring anticyclones remain cold-surface features, 
Johnson (1948) found that typically the Hudson Bay High is a deep system associated with a 
warm ridge aloft in the vicinity of 90°W. This upper ridge is, in turn, related to a trough down 
over western Canada and the United States near 112°W, while the trough off the east coast lies 
near its normal spring position (about 55°W over northern Newfoundland curving gently 
southwestwards). Such situations, which can persist up to five days or more, are preceded by 
blocking over the North Atlantic and western Europe (see also Treidl et al. 1981). In addition, 
Johnson noted that the trajectories of the highs shifted during the course of the season from 
southeastward over New England in March and April to a more southerly route in May. 


Thus the map evidence strongly suggests that the atmospheric circulation over eastern Canada and 
the northeastern United States in June and the first half of July was abnormal in its timing rather 
than in kind - that spring had been delayed or protracted by as much as six weeks. This is in 
keeping with the evidence of the regional climate on Hudson/James Bay, where the normally brief 
summer was essentially obliterated. It points yet again to the importance, climatically, of the 
extraordinary ice and snow conditions over the Bay and northeastern Canada in summer 1816, 
and more generally to the key role of Hudson Bay at this season in the climate of this part of 
North America. 

Concluding Remarks 

For the east coast of Hudson/James Bay during 1809-20, the evidence associating the extreme 
cold with the eruption of Tambora remains circumstantial and intriguing. It is tempting to see the 
period of unrelieved record cold from October 1815 to March 1818 as influenced by high levels 
of stratospheric dust and acid - a relatively short-term feature exacerbating a longer-term 
(60-year) lowering of temperature, which had already begun. There can be no doubt that in 1816 
and 1817 the combination of exceptionally heavy ice lingering through the summer in this part 
of the Bay and the high frequency of onshore north and west winds was a major factor in 
reducing coastal air temperature through the "warm" season. 

From a different perspective, the daily surface weather maps reconstructed for east/central 
Canada, from 1 June to mid-July, indicate circulation patterns and sequences normally related to 
ice and snowcover over the Bay and northeastern Canada, and suggest that spring was running 
some six weeks late in 1816. 

The riches offered by the HBC archives in terms of Canada’s historical climate have scarcely 
been tapped. A major inhibiting factor is the labour-intensive, time-consuming nature of the 
work. The great challenge is to reduce the time and labour required without sacrificing 
information, quality control or physical reality - truly a worthy challenge to modern computer 


Warmest thanks to Gordon McKay, Howard Ferguson and Mal Berry of the Canadian Climate 
Centre for long-term support since 1977 through a series of contracts, which have made the bulk 
of this work possible; also to my Scientific Officers Bruce Findlay and Joan Masterton, to Valerie 
Moore for processing all the manuscripts, to draftsman Brian Taylor (cf. Figures 4,6,7,8), and 
to the many others at the Atmospheric Environment Service who have helped me over these 

I am grateful to Dick Harington of the National Museum of Natural Sciences (now Canadian 
Museum of Nature) for bringing me into the Museum’s Climate Change in Canada Project, and 
for his encouragement particularly with the synoptic-mapping study, for which the Museum 
provided seed moneys. I appreciate too the financial aid and the other help that he has provided, 
with Gail Rice, in publishing my papers in Sy//ogeus, and the work of Edward Hearn (Ottawa 
University) who has drafted nearly all my figures for Syllogeus, most under contract to the 
National Museum. 


I also thank the Hudson’s Bay Company for permission to use the Company archives, HBC 
archivists Joan Craig and Shirlee Smith and their staff, and Alan Cooke, who in 1965 as a 
colleague at the Centre d’études nordiques, Université Laval introduced me to the climatic 
material contained in these archives, thereby opening up a new world. 


Ashmore, S.E. 1952. Records of snowfall in Britain. Quarterly Journal of the Royal 
Meteorological Society 78:629-632. 

Cooke, A. and C. Holland. 1978. The Exploration of Northern Canada, 500 to 1920, A 
Chronology. The Arctic History Press, Toronto. 549 pp. 

Eddy, J.A. 1976. The Maunder Minimum. Science 192:1189-1202. 

Geiger, R. 1965. The Climate Near the Ground. Harvard University Press, Cambridge, 
Massachusetts. 611 pp. 

Johnson, C.B. 1948. Anticyclogenesis in eastern Canada during spring. Bulletin of the American 
Meteorological Society 29:47-55. 

Kington, J. 1988. The Weather of the 1780s over Europe. Cambridge University Press. 166 pp. 
Lamb, H.H. 1988. The weather of 1588 and the Spanish Armada. Weather 43:386-395. 

Landsberg, H. 1967. Physical Climatology. Third edition. Gray Printing Company. Dubois, 
Pennsylvania. 446 pp. 

Manley, G. 1978. Variations in the frequency of snowfall in east-central Scotland, 1708-1975. 
Meteorological Magazine 107:1-16. 

Parker, M.L., L.A. Jozsa, S.G. Johnson and P.A. Bramhall. 1981. Dendrochronological studies 
on the coasts of James Bay and Hudson Bay. Jn: Climatic Change in Canada 2. C.R. 
Harington (ed.). Syllogeus 33:129-188. 

Pfister, C. 1980. The Little Ice Age: thermal and wetness indices for central Europe. Journal of 
Interdisciplinary History 10:665-696. 

Treidl, R.A., E.C. Birch, and P. Sajecki. 1981. Blocking action in the northern hemisphere: a 
climatological study. Atmosphere-Ocean 19:1-23. 

Wilson, C. 1982. The summer season along the east coast of Hudson Bay during the nineteenth 
century. Part I. General introduction; climatic controls; calibration of the instrumental 
temperature data, 1814 to 1821. Canadian Climate Centre Report No. 82-4:1-223. 

. 1983a. Part II. The Little Ice Age on eastern Hudson Bay; summers at Great Whale, Fort 
George, Eastmain, 1814-1821. Canadian Climate Centre Report No. 83-9:1-145. 


. 1983b. Some aspects of the calibration of early Canadian temperature records in the 
Hudson’s Bay Company Archives: a case study for the summer season, eastern 
Hudson/James Bay, 1814 to 1821. In: Climatic Change in Canada 3. C.R. Harington 
(ed.). Syllogeus 49:144-202. 

.1985a. The summer season along the east coast of Hudson Bay during the nineteenth 

century. Part III. Summer thermal and wetness indices. A. Methodology. Canadian 
Climate Centre Report No. 85-3:1-38. 

. 1985b. The Little Ice Age on eastern Hudson/James Bay: the summer weather and climate 
at Great Whale, Fort George and Eastmain, 1814-1821, as derived from the Hudson’s 

Bay Company Records. In: Climatic Change in Canada 5. C.R. Harington (ed.). 
Syllogeus 55:147-190. 

. 1985c. Daily weather maps for Canada, summer 1816 to 1818 - a pilot study. In: 
Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:191-218. 

. 1985d. Daily weather maps for Canada, summer 1816 to 1818. Weather 40:134-140. 

. 1988. The summer season along the east coast of Hudson Bay during the nineteenth 
century. Part III. Summer thermal and wetness indices. B. The indices 1800 to 1900. 
Canadian Climate Centre Report No. 88-3:1-42. 

Wilson, C. and M.A. MacFarlane. 1986. The break-up of Arctic pack ice in 1816 and 1817. 
Weather 41:30-31. 


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Climatic Change, Droughts and Their Social Impact: Central Canada, 
1811-20, a Classic Example 

Dr. Timothy Ball’ 


Changes in the climate of central Canada from 1760 to 1800 were marked by extreme fluctuations 
as the region began to emerge from the nadir of the Little Ice Age. The harshness of climate, 
particularly along the northern limit of trees, created severe ecological conditions. Evidence from 
the historical and meteorological records, maintained primarily by the Hudson’s Bay Company, 
provides clear indications of the extremes and the impact climatic changes had on the socio- 
economic infrastructure of the region. Between 1800 and 1810 the climate was relatively benign, 
holding promise of better conditions and times in the nineteenth century. The promise was short- 
lived as temperatures began to decline in 1811, a trend that was to continue through to 1818. 
Most research on the period has stressed the temperatures, but detailed studies of the historical 
documents show that the period from 1815 to 1819 was one of severe drought as well as cold. 
The combination suggests that the mechanisms causing the drought were probably different than 
those that created the hot droughts of the 1930s. 

Droughts in the Canadian prairies are usually attributed to a northward extension of the Pacific 
High (Subtropical). Droughts in the boreal forest region are usually associated with a southern 
position of the Arctic High. Both regions indicate a 22-year cycle of droughts that seems to 
coincide with sunspot cycles. With a northward shift of the mean summer position of the Arctic 
(Polar) Front there is a hot drought on the prairies. With a southerly location of the Front there 
is a cold drought. The 1815-17 period was a classic example of the latter, and is associated with 
an extreme degree of meridionality in the zonal index. 

Variations in precipitation in the early nineteenth century had a significant effect upon the wildlife 
of the region. This resulted in a decrease in the food supply for Europeans and Indians, with the 
concomitant social and health stresses. A decline in the fur-bearing animals created declines in 
income that led to significant socio-economic adjustments. 


A great deal of attention has been paid to the exceptionally cold summer of 1816. A study of the 
eventful summer has been documented by Hoyt (1958) with a description of weather, food 
supplies, prices, and even population movements. Work focused upon the northeastern United 
States and Canada. Attention spread to what Post (1977) called, "the last great subsistence crisis 
of the western world". The most extensive and analytical study was the book Volcano Weather 
(Stommel and Stommel 1983). This work included a chapter on conditions reported in Europe. 
A brief reference at the end of the chapter suggests limits to the extent of the area influenced by 
cold conditions, "It would appear, however, that the truly exceptional character of 1816 weather 
was limited to a small portion of northeastern America, Canada and the extreme western parts 
of Europe" (Stommel and Stommel 1983, p. 51). 

! Department of Geography, University of Winnipeg, 515 Portage Avenue, Winnipeg, Manitoba R3B 2E9, Canada. 


The severe cold conditions have more recently been detailed in the Hudson Bay region of 
northern Canada (Catchpole 1985; Catchpole and Faurer 1985; Skinner 1985; and Wilson 1985). 
Such work has produced not only valuable information on the extent and intensity of the cold 
conditions in 1816, but it has also contributed to ideas concerning the possible causes of such 
exceptional conditions. 

The debate centres on whether the cold conditions were caused by: (1) the intensity of the dust 
veil emitted by the equatorial eruption of Mount Tambora in April 1815; or (2) the influence of 
variable sunspot activity; or (3) natural variation caused by some atmospheric, or 
atmosphere/ocean phenomenon. This controversy remains unresolved. While each effect may 
exert an influence on large-scale atmospheric circulation, by reducing the Earth’s radiation 
balance, they may also exert a simultaneous effect. 

Documentary evidence for the summer of 1816 appears to agree on one aspect of the cold 
summer - that a ridge of high pressure extended south over eastern North America and western 
Europe bringing cold arctic air well south of its normal latitudes for the time of year. These 
systems are relatively common in the fall, winter and early spring, but are unusual in the 
summer. Their impact on the socio-economic conditions of that period were severe. 

Here, I intend to show that the pattern of weather in 1816 can be generally defined from climatic 
information in Hudson’s Bay Company records. The pattern indicates that cold conditions did not 
include the entire prairie region. Southern Alberta had normal conditions, while the north had 
an exceptionally wet summer. Overall weather conditions began to deteriorate in 1809, and 
continued to decline until 1816. That year, apart from being cold, was the first year of a severe 
drought that lasted until 1819. Comparison of conditions with modern synoptic charts suggest that 
this was a cold drought within the 22-year cycle of droughts experienced in the Great Plains. 

The 22-year cycle of droughts correlating with sunspot activity has generally been established and 
accepted (Herman and Goldberg 1978). Very little detailed analysis of the nature of each drought 
period has been completed. It is generally accepted that droughts are coincident with hot weather, 
and the 1930s are cited as the classic example. Undoubtedly hot dry weather is especially 
damaging to modern agriculture, but lack of precipitation under any temperature regime is 
serious. The cold drought from 1816 to 1819 was especially damaging, as journals and diaries 

These years of severe weather had a considerable impact upon wildlife, indigenous peoples and 
Europeans. Later, I will suggest that it served as a catalyst for an already volatile situation: the 
Seven Oaks massacre in 1817 at the Red River settlement. 

The fur trade had been suffering from over-trapping and competition between the North West 
Company and the Hudson’s Bay Company. Tensions between Indians, Métis and fur traders were 
somewhat overshadowed by the growing talk of permanent European settlement. The first group 
to arrive, the Selkirk Settlers, came from Scotland in 1811. Generally, they were unwelcome 
because they threatened the fur trade and the traditional ways of native people. Ironically, they 
had left Scotland because of severe weather. Now they had moved into a land that was suffering 
for the same reason. Diminished wildlife populations meant reduced food supply, with associated 
hunger and disease. Residents already felt threatened and unsure, thus they saw the settlers as an 
even greater threat. This situation heightened tensions and began a long period of conflict. 


The Hudson’s Bay Company Post Journals often provide a brief but daily summary of activities 
and weather conditions. Some Journals are incomplete in this period because of feuds between 
the Hudson’s Bay and the North West companies, or absences from the Posts on expeditions 
between Inland Posts and Bay Posts to exchange furs for supplies. Sometimes the mere struggle 
for survival precluded maintenance of the records. Although the Journals are fragmented, the 
entries provide a series of proxy data that give some indication of the activities and weather 
conditions that occurred within these years. 

Data Sources 

Proxy data from the Journals include: (1) comments on garden-crop preparation, growth and 
damage; (2) remarks of frost or ice formation, movement and decay; (3) descriptive terms for 
winds or precipitation restricting outdoor activities or travel. Even limited comments of 
phenological data for animal appearances or migrations serve as an indicator of weather 

The Hudson’s Bay Company Posts located in the west-central region of Canada are shown in 
Figure 1. Brandon House was an important Post adjacent to the Assiniboine River. Peter Fidler, 
the Factor in charge, provided much of the proxy data and insights into the weather through his 
records. Carlton House, although shifted several times, was transferred in 1810 to a site near "a 
crossing place" on the south bank of the North Saskatchewan River. The daily accounts written 
for each Post between May and October 1810 to 1820, were analyzed for proxy data that might 
be attributable to adverse weather. 



Simpson , 


a Kaniapiskau 

Nichikunae — 



Hauge Mistassini 

Carlton, Moose 

Lac Osnaburgh 
Red Seul Ld rm 
. Be s Nipigon 
Brandon House® . 

* Waswanipi 


g Matawagamingue 

oO 400 800 

Figure 1: Location of Hudson’s Bay Company posts. 


Gardening and crop production were a major part of the general way of life. The Company 
encouraged each Post to obtain much of its sustenance from local sources in order to be as self- 
sufficient as possible. Gardens were maintained, and hunting and fishing supplemented the diet 
with a fresh supply of meat. Cutting firewood for the approaching winter also took up much of 
the time (Ball 1987). Frequent canoe trips to Posts along Hudson Bay were embarked upon in 
the short summer season. The observations recorded during these activities are evidence of the 
severe cold experienced in the early to late summers of 1816 and 1817. Warmer weather returned 
in 1818, but the drought continued to 1819. 

Peter Fidler’s daily records for Brandon House yield valuable information. In 1816, the ice must 
have dispersed toward the end of April or early May as Indians were fencing in the Assiniboine 
River on 21 May "...half of mile above the House to kill sturgeon" (HBCA, PAM B22/a/19).: 
Fidler mentioned that the presence of sturgeon usually gives an indication of when the ice goes 
out "...they annually come up every spring in great numbers when the ice goes away and they 
appear here about 10 to 12 days after it clears away..." (HBCA, PAM B22/a/19)'. 

Gardening and crop preparation began on 3 May, and an indication of the spring runoff levels 
and weather were observed by Fidler on 9 May as water levels were "...falling daily 1% inch - 
cold weather and strong wind these two days" (HBCA, PAM B22/a/19). The remainder of the 
crops were sown by 27 May, but it was not until 5 June that a severe cold spell occurred. "A 
very sharp frost at night and killed all the Barley, Wheat, Oats and garden stuff above the ground 
except lettuce and onions - the Oak leaves just coming out are as if they are singed by fire and 
dead" (HBCA, PAM B22/a/19). With a severe frost early in June the growth of crops and natural 
vegetation would certainly be curtailed, as this period is essential for their development to mature 

An interruption in the Journal occurs after this period due to the battle between the Hudson’s Bay 
Company and the North West Company’. In the spring of 1817, Brandon House was subjected 
to severe weather, as on 18 June "...thin snow fell 2 inches deep" (HBCA, PAM B22/a/20). 
Visitors were late arriving at the Post due to the backward spring, and Fidler explained the cause 
of late arrival; "...he was detained long by the ice in the Little Winnipeg" (HBCA, PAM 
B22/a/20). Little Winnipeg refers to Lake Winnipegosis. The summer of 1817 was reported to 
be backward due to the unseasonable cold, but drought also had a direct effect on agriculture. 
"The crops exceedingly backwards - some potatoes only 4 inches above ground - whereas in other 
seasons there were new ones bigger than Walnuts, the grass is also remarkably short and ground 
dry - all the little runs of water now dry - so there is every reason to expect a bad crop on 
account of the great want of rain - the season has been colder than usual" (HBCA, PAM 

The summer continued to be dry, as the small saline lakes began to evaporate. The apparent 
migrations of buffalo southward also give an indication of the dry summer and unusual cold. 
Fidler recorded buffalo movements near the Post in the spring, and by 11 August they were 
"...very numerous - even extending so low down as the Forks" (HBCA, PAM B22/a/20). The 
cold weather continued. Fidler wrote in August of frosts that occurred on 17 and 23 July at 
Brandon which killed all the potato tops. The autumn season seems to end on 23 October when 

' Hudson’s Bay Company Archives, (HBCA) Provincial Archives of Manitoba (PAM), Journal Number. 
? The conflict between the two companies was resolved in 1821 when the smaller Hudson’s Bay Company 
incorporated the North West Company. 


the Assiniboine River froze over. This occurrence is seen by Fidler as being "...very early in the 
season, about 20 days sooner than usual - and it set in early last fall" (HBCA, PAM, B22/a/20). 

The spring of 1818 apparently began without mention of adverse conditions, as Fidler reports on 
26 May "...the ice drove by about 5 weeks ago ..." (HBCA, PAM, B22/a/21). Despite a break 
in the daily reports for Brandon House, Fidler continued to write on his journey from Red River 
to Martins Falls near Albany Factory on James Bay. 

Returning to Brandon House, Fidler recorded the late summer conditions: "Water very low in 
the river and a very dry season scarce a single shower of rain all summer, all the potatoes and 
garden stuff quite burnt out as also 2% bushels of Barley sown there - when 3 inches high all 
killed by the great drought - these 3 summer past remarkably little rain ... quite different from 
what it used to be" (HBCA, PAM, B22/a/21). 

Summers at Carlton House are also recorded in fragments due to continuing battles between 
Hudson’s Bay and North West companies, and canoe trips to other Posts. However, direct 
information recorded in the Journals still provides an indication of summer conditions for 

The spring of 1815 appeared to have a positive beginning: crops were planted as early as 
29 April. However, conditions changed, and on 13 May John Pruden recorded the bleakness of 
the weather; "...hard frosts every night retards vegetation very much, none of the seeds that have 
been sown make their appearance above ground except the cabbage seed" (HBCA, PAM, 

By June, weather continued poor, but there was a different problem. "Wind SW blowing fresh 
part-clear and part cloudy weather, it has been remarkable dry wind weather all this month which 
keeps the garden stuff very backward" (HBCA, PAM, B27/a/5,2d). Things had not improved a 
month later: "The insects have eaten all our cabbage and turnips owing I suppose to the dry 
season, ..." (HBCA, PAM, B27/a/5,4). This was the first indication of a drought that was to grip 
the eastern half of the prairies for three years. The impact was to be quite severe. 

The drought conditions are best summarized in Peter Fidler’s General Report of the Red River 
District for 1819. 

The spring months have sometimes storms of wind and thunder even so early as 
March within these last years the Climate seems to be greatly changed the 
summers so backward with very little rain and even snow in winter much less than 
usual and the ground parched up that all summer have entirely dried up, for these 
several years loaded craft could ascend up as high as the Elbow or Carlton House 
but these last 3 summers it was necessary to convey all the goods from the Forks 
by land in Carts... (HBCA, PAM, B22/e/1,6). 

We can discover the extent of the drought by noting which rivers are reported to be low. The 
North Saskatchewan, Assiniboine, Red, Hayes, Nelson and Steel rivers all receive attention in 
the journals. This means that the drought covered the drainage basins of all of these rivers, thus 
encompassing a large part of central North America. 

The degree of the drought can be determined by the impact that it had on the environment, 
wildlife and subsequently the people. James Sutherland reports that water routes connecting the 


Hayes and Nelson rivers were only made passable by the construction of dams (HBCA, PAM, 
B154/e/1,2). Peter Fidler notes that: 

...aS the country wherever I have been and by the invariable information of the 
different Tribes I have enquired at agree the country is becoming much drier than 
formerly and numbers of small lakes become good firm land will be covered with 
Timber of various kinds...CHBCA, PAM, B22/e/1,8d). 

Fidler implies that he expects these conditions to persist in the future, although he does not 
specify for how long. 

The value of his comments lie in putting the individual events into a larger and longer climatic 
framework. It is important to note that all seasons suffered from the lack of precipitation. "These 
3 summers past remarkably little rain - as also very little snow in winter quite different from what 
it used to be" (HBCA, PAM, B22/a/21,29d). We also know that conditions were good prior to 
the drought; "...since 1812 there was always good crops of everything until 1816 when the dry 
summers commenced..." (HBCA, PAM, B22/e/1,8). 

Prairie droughts are usually accompanied by the appearance of insects, particularly grasshoppers, 
that exacerbate the problems. Fidler makes some interesting comments when talking about the 
grasshopper infestation. He notes that, "...They first made their appearance the third week of 
August 1818 at 2 O’clock in the afternoon and came from the southwest" (HCBA, PAM, 
B22/e/1,20). The direction is significant because it indicates wind direction during the period. 
This is confirmed by John Pruden’s observation at Carlton House that, "Wind SW blowing fresh 
part-clear and part cloudy weather, it has been remarkable dry wind weather all this month..." 
(HBCA, PAM, B27/a/5,2d). Then Fidler writes: "These insects (grass-hoppers) make their 
appearances in great numbers about every 18 years..." (HBCA, PAM, B22/e/1,6d). This implies 
cycles of infestation and possibly of climate. 

Atmospheric Circulation 

So far we have established that 1811-20 had below normal temperatures, especially in the years 
1816 and 1817. Between 1812 and 1816, conditions were cool but generally good for crops and 
vegetables. In 1816 drought began in a large region including the central and eastern prairies. 
The drought ended in 1820 as temperatures and precipitation patterns returned to long-term 

How did the circulation pattern for these years differ from the long-term normal, and was the 
drought typical of those that occur regularly on the prairies? 

The climate of central Canada is generally determined by the position of the Arctic Front’. In 
summer the mean position of the Front approximates the northern boreal forest limit. In winter 
it curves south in a great arc toward the centre of the continent to an approximate mean position 

' There appears to be some confusion over the use of the term Arctic Front. Bryson and others have used the term 
Arctic Front to describe the major division between Arctic and Temperate air in North America. The term Polar 
Front is used by others, particularly in Europe, presumably to indicate that there is a similar front in the southern 
hemisphere. I have used Arctic Front because the paper is examining conditions in North America. 


of 40°N Latitude. This curve occurs because the Rocky Mountains act as a barrier, and create 
a standing wave in the westerly flow of the general circulation. 

Polar air north of the Arctic Front tends to be cold and dry, while subtropical air to the south 
tends to be warm and dry. Generally, moisture is brought to the region by cyclonic storms that 
move along the front; these usually occur in spring and autumn as the Front moves through the 
region in its annual migration. Most summer precipitation is convectional as instability develops 
in the warm subtropical air. 

Dey (1973) analyzed synoptic conditions occurring during summer dry spells in the Canadian 
Prairies. He showed that the most severe droughts occurred when the Pacific High (subtropical) 
extended northward into the southern prairie region. Blocking occurs with the extreme meridional 
pattern that is formed. This configuration is commonly called an ‘omega block’. Low pressure 
zones on the Pacific coast and in the region to the west of Lake Superior lie on each side of a 
large high pressure region. On the weather map this creates a pattern similar to the Greek letter 
omega - hence the name. 

It is also possible to have dry conditions in summer if the Arctic Front extends southward over 
the region. This would produce cool, dry conditions under a predominantly northerly flow. The 
summers of 1815-17 are good examples, being marked by very cool dry conditions in the early 
summer as the Arctic Front remains well south of its normal position. When the Front finally 
retreats northward, the prairie sites experience the associated precipitation. For example, at 
Carlton House in July 1815, the Journal reads: "Wind easterly cloudy weather had a heavy 
shower of rain last night, the only one I may say since summer commenced..." (HBCA, PAM, 
B27/a/5,3). In June 1817, Swan River experienced three days of continual rain under cyclonic 
conditions as the Front migrated northward. 

Further evidence to support this hypothesis is provided by the weather patterns at Fort 
Chipewyan. This post is ideally located to determine the latitudinal and longitudinal shifts of the 
Arctic Front in summer. Spring was late each year from 1815 to 1818 inclusive. The summers 
were cool and short, as autumn came early. This was especially true in 1816 as the entry for 
30 September indicates: "One of the several days that I have witnessed at this season of the year, 
the ground covered with snow 31% inches deep, blowing very fresh and extremely cold" (HBCA, 
PAM, B39/a/9,10d). 1818 saw the return to a longer summer, with the first snow falling on 
13 October, and a comment that there was "...mild weather with wind from the south" on 
22 October (HBCA, PAM, B39/a/14,8d). 

fle-a-la-Crosse, south of Fort Chipewyan, has a limited record, but it does indicate normal 
conditions. For example, the earliest date of the water being clear of ice in the modern record 
is 12 May. In 1816 the ice was gone by 18 May. It is reasonable to use this station as the western 
limit of the outbreak of cold arctic air in the spring and early summer of that year. 

In summary, it appears that 1815 saw the beginning of generally cooler conditions. The cold was 
more notable in 1816 and 1817, especially for the spring and early summer in the eastern half 
of the Prairies. The Jetstream and Arctic Front swung south so that the eastern half of the Prairies 
was cold and dry, under arctic air. Conditions changed significantly in 1818. The Arctic Front 
moved north, and an omega block system set in that appeared to dominate through 1819. Thus 
the cold drought that existed in 1816 and 1817 was replaced by a warm drought in 1818 and 




45° bh 45° 



120° 902 60° 30° o° 30° 60° 

Figure 2: General reconstruction of the pressure patterns for North America and the North Atlantic for 
July of 1816 (after Catchpole 1985; Lamb and Johnson 1966). 

> 30 
rising, snow 

: 29 WV 
i (43) Alightly Overcast 

(No geese flying) ~ 
ea te ee 

. ° 
very fine morning 

cold | 
severe frost 
snow last night 
: ! very cold, snowing) ¢ y i —§$———— 
__ A very disagreable night, oo TOSthes y @ | ( (40) 
a very sharp frost (snow ont 
(killed plants, singed oak leaves) : ground) 
Te oe : very hard frosté 
iene i \cold}* 

ery clearg i 
‘...(snow on ground) 8% cloudy, snowne 

32 : 
(33) : 

snowing \ 



: cold, unpleasant day 
June 5 1816 (rain began 2p.m.) 
Morning se 

(For legend see Figure 5a.) 
———~. Schematic pressure patterns 

Fronts: — warm, =< cold. 

° 200 400 800 800 1000 

ec ot rr 1 OME TRES 


Figure 3: Surface weather map: morning, 5 June 1816 (after Wilson 1985). 


— Surface isobars 

-—-—- 700mb contour 

s+eeee* 500 mb contour 
é ij Jet stream 


Figure 4a: Synoptic weather pattern for drought conditions in Western Canada (after Dey 1973). 

Figure 4b: Synoptic weather pattern for drought conditions in Western Canada (after Dey 1973). 


This pattern is consistent with the general reconstruction suggested by Catchpole (1985) 
(Figure 2). It also demonstrates that Wilson’s (1985) diligently drawn synoptic maps for three 
months of the summer of 1816 are valid. Wilson’s map for 5 June 1816 (Figure 3) shows the 
surface conditions with a southern expansion of the Arctic High. Synoptic conditions as 
reconstructed by Dey (1973) would have been the general situation in 1818 and 1819 (Figure 4). 

Climatic Impact 

Regardless of the climatic mechanisms, there is no doubt that climatic conditions seriously 
affected wildlife and people’s ability to grow food. The lack of precipitation reduced the planted 
crops, but it also affected the wild harvest of berries and other fruits. Lack of snow is alas 
to many wildlife species, especially with colder-than-usual temperatures. 

Climate also affected people’s ability to travel. Early ice, late ice, too little snow, as well as 
shallow rivers and lakes all hampered movement for trade and hunting. Buffalo migration also 
indicates unusual conditions. Normally these animals would move westward or southward, but 
with changing snow patterns and dry conditions to the west they altered their behaviour. Fidler 
writes in 1817: "There are plenty of Buffalo not 15 miles off and all last winter and this spring 
they have been very numerous - extending even so low down as the Forks" (HBCA, PAM, 
B22/a/20,9d). Again in 1818: "The Catfish during these summers have also been very 
scarce...but fortunately vast numbers of Buffalo have kept pretty near all summer" (HBCA, 
PAM, B22/a/21,2). 

Life has always been a struggle in this part of the world. Food supply varies dramatically with 
climate. It is truly a region of plenty or dearth. However, the 1811-20 period was one of special 
severity. The 1780s and 1790s had been periods of severe climatic conditions, with weather 
oscillating from one extreme to another. The author has argued elsewhere that this period created 
the pressure that forced the Company away from its complacent position on the Bay. A brief 
respite from 1800 to 1810 was then shattered by severe cold and drought. Starvation and hardship 
returned, as animals disappeared or changed their routines. Fur traders saw their industry 
threatened. Indians saw the fur trade threatened and their traditional way of life dashed. Conflict 
for the trade was at a peak as the Hudson’s Bay and North West companies literally battled over 
the spoils. The Selkirk Settlers entered the scene unaware of the tensions and problems. Since 
the settlers posed a threat to all, it is not surprising that a series of confrontations occurred 
culminating in the massacre at Seven Oaks in 1817, when 20 men were killed. The stress of the 
impending social and cultural confrontation was underlain by exceptionally severe weather. 
Starvation and malnutrition made rational behaviour less likely. 


There appear to be two types of drought on the Great Plains of North America; hot droughts and 
cold droughts. The former coincide with the omega blocking system that dominated the region 
in 1988. In this system the Pacific High extends northward over southern Alberta, southern and 
central Saskatchewan, and southern Manitoba. This creates very hot and dry conditions similar 
to those seen in the 1930s and again in the 1980s. The latter occur when the Arctic Front dips 
south across these same regions so that they are now dominated by clear, cool and equally dry 


The 1816-19 period was one in which cold drought predominated. Journals of the Hudson’s Bay 
Company provide much information about the extent and intensity of the conditions. They also 
allow estimation of the impact that these conditions had upon the environment, and therefore upon 
wildlife and people. 

The heat and drought of the 1980s have led to current predictions of global warming and 
impending doom as droughts increase in frequency and severity in North America. My brief 
study suggests that this will not be the case. Perhaps the pattern of hot or cold droughts will 
change. A more northerly location of the Arctic Front might result in less southerly incursions 
of Arctic air. 


Ball, T.F. 1987. Timber! Beaver 67(2):45-56. 

Catchpole, A.J.W. 1985. Evidence from Hudson Bay Region of severe cold in the summer of 
1816. In: Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:121-146. 

Catchpole, A.J.W. and M.-A. Faurer. 1985. Ships’ Log-Books, sea ice and the cold summer of 
1816 in Hudson Bay and its approaches. Arctic 38:(2)121-128. 

Dey, B. 1973. Synoptic climatological aspects of summer dry spells in the Canadian Prairies. 
Unpublished Ph.D. thesis. University of Saskatchewan. Saskatoon. 180 pp. 

Herman, J.R. and R.A. Goldberg. 1978. Sun, Weather and Climate. Scientific and Technical 
Information Office, NASA, Washington, D.C. Sp-426. 360 pp. 

Hoyt, J.B. 1958. The cold summer of 1816. Annals of the American Association of Geographers 

Lamb, H.H. and A.I. Johnson. 1966. Secular variations of the atmospheric circulation since 
1750. Geophysical Memoirs 110. H.M.S.O., London. 125 pp. 

Post, J.D. 1977. The Last Great Subsistence Crisis in the Western World. Johns Hopkins 
University Press, Baltimore. 240 pp. 

Skinner, W.R. 1985. The effects of major volcanic eruptions on Canadian climate. Jn: Climatic 
Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:75-106. 

Stommel, H. and E. Stommel. 1983. Volcano Weather. Seven Seas Press Inc., Newport, Rhode 
Island. 177 pp. 

Wilson, C.V. 1985. Daily weather maps for Canada, summers 1816 to 1818 - a pilot study. Jn: 
Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:191-218. 


The Year without a Summer: Its Impact on the Fur Trade and History 
of Western Canada 

Timothy F. Ball’? 


Edward Umfreville referred to the Hudson’s Bay Company as “being asleep by the frozen sea". 
He was talking about the fact that the Company had established its trading posts along the shores 
of Hudson Bay and made no attempt to build permanent posts inland. Arthur Dobbes used this . 
as evidence in his charge of monopoly against the Company. He claimed that the Company was 
deliberately protecting and hiding the potential of the interior of North America to ensure the 
dominance of the Hudson’s Bay Company. 

At the end of the eighteenth century the Company established its first inland post at Cumberland 
House on the Saskatchewan River. It has always been argued that the sole reason for this move 
was to counteract the expansionism of the North West Company. Increasing evidence suggests 
that climatic change brought about a dramatic decline in the ecology of the northern region and 
this was a major cause of the move inland. By 1810 the expansion created increasing conflict 
between the two companies. In 1812. a third component, the Selkirk Settlers, arrived and the 
turmoil continued to build. 

Severe weather affected all three groups through their dependence upon the land for sustenance 
and economic profit. Two events, the Seven Oaks Massacre and the amalgamation of the 
Hudson’s Bay and North West companies, followed the period of most severe weather in 1816- 
17. There is little doubt that the summer of 1816 was one of the worst in the historic record. It 
was referred to as "the year with no summer" and, more recently, as "the last great subsistence 
crisis in the western world". The effects on the living conditions in Western Canada were well- 
documented, and clearly placed a great deal of stress on native people and the European traders 
and colonists. Friction between the groups was exacerbated by the uncertainties of food supply. 
Probably the hardships created by the extreme weather were a significant catalyst for the events 
that occurred. 


The effect of climate on human behaviour has been a contentious issue in the twentieth century. 
The concepts that evolved from Friedrich Ratzel’s Anthropogeography, published at the end of 
the nineteenth century, were transported and transposed by various people until they came to a 
distorted rest in Adolf Hitler’s Mein Kampf. Since then the concept of climate influencing people 
or history has been anathema in the academic world. Unfortunately, it is evident from even a 
cursory glance at the patterns of climate and the sequence of history that we “threw the baby out 
with the bathwater’. 

' Department of Geography, University of Winnipeg, 515 Portage Avenue, Winnipeg, Manitoba R3B 2E9, Canada. 

2 The following article is a précis of a public lecture given during the conference The Year Without a Summer? 

Climate in 1816 at the National Museum of Natural Sciences. 


The purpose of this presentation is not to pursue the idea of climatic determinism, but rather to 
examine the pattern of the fur trade in the context of climatic conditions. The argument is 
presented that plants and subsequently all animals are limited in their options and reactions by 
climatic conditions. History must be examined in the context of climate because of its control 
over the fundamentals of life. I prefer to think that geography and history are inseparable; history 
is the play and geography the stage on which it is enacted. 

It is interesting that anthropologists have little problem with the idea that primitive societies are 
essentially controlled by climate, but somehow historians reject the idea. What is the difference 
between the two? It is partly the fact that, until relatively recently, we have known little about 
the climate of this historic period. There is also a great deal of conceit in the belief that humans 
are not as affected by climate as other animals. This conceit has reached its highest levels in 
North America in the twentieth century where technology is believed to have the answers to all 
problems. Despite the fact that climate dictates over 80% of the yield on any farm there are no 
compulsory courses in climate or meteorology at Canadian schools of agriculture. The drought 
of 1988 brought the realities of the dominance of climate to the fore once again. It should have 
reminded us that man’s mastery over the environment is a figment of his conceited imagination. 
I hope that, as we increase the amount of knowledge about past climates, we include it as a 
significant factor in the pattern of human actions; both past, present and future. 

The pioneering work of people like Hubert Lamb delving into historical diaries and journals to 
reveal very different past climatic conditions has only occurred since the Second World War. 
Historical climatology has shown that climate has varied a great deal in time and space, thus 
altering the prosperity of different regions. The primary alteration is in the ability to produce 
food. However, climate also affects commerce, especially if it depends upon a natural product 
that is weather dependent. 

One cannot examine the impact of the period from 1789 to 1820 upon the fur trade without 
considering the broader context. Rarely do singular climatic periods or events create direct 
change. Invariably a system is put under increasing pressure until certain climatic conditions 
become a catalyst to change. 

The fur trade in North America is a good example of an enterprise almost totally dependant upon 
climate for its survival and success. Climate dictates: the number and quality of furs; conditions 
for the trappers and their families; the ease of transport through snow conditions cr water levels 
in rivers and lakes; the ease of shipment across the oceans; the dependency of Europeans upon 
food supply from the land, to name a few items. 

There is no point in blaming historians for ignoring climate as a factor in such change because 
the information has not been available. As reconstruction of climatic patterns continues, it is 
essential this be included as a major factor in the mosaic of variables that direct the human 

The period from 950 to 1200 is variously referred to as the medieval warm epoch or the Little 
Climatic Optimum. Regardless of the term, it was a period of much warmer conditions than at 
present. Oats and barley were grown in Iceland; the Domesday Book records commercial 
vineyards flourishing in England; while the eleventh and twelfth centuries later became known 
as the golden age in Scotland. 


It is important to note what was happening in North America because of the parallels with current 
predictions of global warming. The warmer conditions resulted in northward migration of the 
agricultural people of the lower Mississippi Valley into Wisconsin and Minnesota. However, it 
also resulted in increasing aridity in the Midwest - that is the area west of the Mississippi. 
Analyses of Holocene pollen from the northern plains of Iowa indicate increasing aridity and a 
change from deciduous forest to grasslands. In Canada the northern limit of trees expanded 
northward up to 100 km in some regions, as warmer conditions brought a longer growing season. 

After 1200, global climate began to cool. The circumpolar vortex expanded and the zone of 
cyclonic storm activity shifted south. Cultures that had benefited from the warmer conditions now 
saw a decline, but as with any climatic shift, others gained. For example, the Old Norse colony 
in Greenland collapsed as crop failures increased and permafrost returned. The settlements in - 
Iceland and Norway experienced a decline in population as agricultural conditions deteriorated. 
In North America the increased strength of the westerlies resulted in a greater rainshadow effect 
in the lee of the Rocky Mountains and increased dryness on the Great Plains. 

The problems in Europe were a litany of woes for people who had experienced the warm 
conditions of the Little Climatic Optimum. The woes included: increasing storm severity; harvest 
failures; abandonment of croplands and villages in higher elevations; and an increase in disease 
and mortality rates. In Scotland it has been estimated that the elevation at which agriculture could 
be practised lowered by 200 m between 1450 and 1600. The greatest loss was in the Highlands 
because the vertical loss converts into a substantial horizontal loss, which is devastating in a 
country with little level or arable land. 

Martin Parry has estimated that harvest failures occurred one year in 20 in the thirteenth century. 
By the late seventeenth century this had been reduced to one year in two. Consecutive years of 
failure led to consumption of seed grain, thus accentuating the situation. The implications are that 
the initial Highland clearances were caused by climate, not by land-hungry landlords. With 
Highland clans forced to lower ground, the clan wars began. This was to be the beginning of 
many decades of extreme social upheaval. In 1675 the Philosophical Transactions of the Royal 
Society reported that a lake in Strathglass had " on it in the middle, even in the hottest 
summer." It is also reported by others that there was permanent snow on the tops of the 
Cairngorms. We rarely stop to think that curling, which originated on lake ice, could not be 
played in many winters in the twentieth century. 

It has been estimated that by 1691 over 100,000 Scots had been transplanted to Ulster, driven by 
such conditions as occurred in March 1674 when excessively heavy snows, severe frosts and 
13 days of drifting snow resulted in the deaths of hundreds of sheep. Unfortunately the 1690s 
brought even worse conditions; in the eight years from 1693 to 1700 there were seven failures 
of the essential oats harvest. An excellent measure of the degree of cold during this general 
period was the winter of 1683, known as the year of the great frost, when 2 feet (0.6 m) of ice 
formed on the Thames River in London. 

Hubert Lamb called 1450 to 1850 the Little Ice Age. The coldest portion of this period was from 
1645 to 1715, with the nadir occurring in the 1690s. (Ironically this period coincides closely with 
the lifespan of the astronomer, Edmund Halley). The impact of these climatic conditions on 
Europe are receiving growing attention. A very important point is that this 70-year period 
coincides with a period known as the Maunder Minimum. During this time there were virtually 
no sunspots. There is increasing evidence that when there are high numbers of sunspots, as in 
the 1980s, the Earth is warm, and when there are few it is cold. 


Although the initial hardships of the Little Ice Age were caused by deteriorating weather, the 
reaction of the landowners was, in most cases, reprehensible, ranging from absolute cruelty to 
benign neglect. One who did attempt to alleviate the problems faced by some of his tenants was 
Thomas Douglas, 5th Earl of Selkirk. It is essential to understand that even his motives are 
suspect. That is they were not as altruistic as people have proposed. His objective was to ensure 
colonization of the North American continent to halt expansion of the American revolutionary 
nation. He detested revolutionaries, but especially Americans. The family house had been 
attacked by John Paul Jones, Scottish-born naval hero of the American War of Independence. 
Young Selkirk - then seven years old - was so frightened and angered that he held a lifelong 
grudge against Americans. 

Climate and Fur Trade in Western Canada 

I will return to Lord Selkirk later. It is necessary now to look at the evolution of the fur trade. 
I believe it is significant that the Hudson’s Bay Company’ received its charter in 1670, just 10 
years before one of the coldest decades in the last several hundred years. The demand for furs 
would have been much different in the warmth of the Little Climatic Optimum. The Company 
prospered as the demand for furs increased, and they expanded their operation accordingly. 
Interestingly they did not move inland from the shores of the Bay remaining, as Edward 
Umfreville described it, "asleep by the frozen sea." This was to place increasing pressure on the 
wildlife in the northern regions, which will be discussed later. 

The first half of the eighteenth century saw a gradual growth of the fur trade. A widening region 
yielded more and more furs, but now there was a growing confrontation. The pedlars or 
Canadians (as the Hudson’s Bay Company called the fur traders of the North West Company 
operating from Quebec) expanded westward across the prairies. Debate began within the 
Company about the need to open inland posts to offset the threat. Historians argue that the 
decision to move inland was totally due to this competitive factor. I contend that competition was 
a factor, but more significant was the impact of climate, especially in the northern regions around 
and west of the Bay. 

While the debate was occurring, global climate was changing. The weather records for Churchill 
and York Factory show a significant shift in the pattern of winds, precipitation and other 
variables. Prior to 1760 the mean summer position of the Polar Front was south of Churchill and 
York Factory. This meant that both sites experienced subarctic climatic conditions. After 1760 
the Front shifted north, so that Churchill continued with a subarctic climate but York Factory 
now had a more temperate climate capable of supporting the boreal forest. By the 1780s another 
shift in climate was occurring. Conditions deteriorated and the record becomes replete with 
comments on the lack of game, hard times and starvation among the Indians. At the same time 
there was a continued reduction in the number of furs being taken. 

A study of the period from 1780 to 1800 by Stephen Wilkerson quantified the deteriorating 
conditions. Content analysis (a frequency count of the number of references to starvation and 
other key words) clearly shows a system under extreme stress. In some years, such as 1792 the 
number of comments about lack of food, starving people, malnutrition, and death are eight to 10 
times above previous periods. An entry for 17 January 1792 reads, "Indeed this winter has been 

'! The body of water is named Hudson Bay: the company is correctly called "The Hudson’s Bay Company" or 

sometimes just "the Company". 


so far the most remarkable for scarcity of provisions for neither Englishman or Indians can find 
anything to kill." 

These conditions are coincident with unusual patterns of animal behaviour. Joseph Colen recorded 
the following in the Journal for 22 October 1787, "Game of all kinds scarce but that White Bears 
are SO numerous and trouble some as to attack them and their stages where their provisions is 
deposited." Later in the same year an entry for 23 December reads, "Late in the evening large 
herd of wolves surrounded the Factory." Both these events are unusual for the period. Colen was 
to become a victim of the conditions. 

During this period the number of furs taken was significantly reduced. Colen wrote to London 
arguing that the cause was overtrapping. The Company accused him of mismanagement and 

removed him from his post. Actually he was right, but for the wrong reason. The number of furs 

taken would not normally have been a problem except that now the climatic changes had altered 

the thresholds. Wide variations in temperature created great stress on the plants and animals. For 

example, an entry for 5 February illustrates the unusual nature of the situation, "They (Indians) 

also inform me that the winter set in so early upwards that many Swans and other waterfowl were 

froze in the Lakes and they found many of the former not fledged, they likewise say that the 

snow is remarkably deep." 

However variations in precipitation caused the greatest difficulties. An entry for 14 November 
1783 reads, "Never remember the snow so deep at this season of the year." The next year an 
entry for 24 April informs that "Snow at least 10 feet deep." Too much snow creates great 
difficulty, especially for larger animals including man. Too little snow is devastating. Ptarmigan, 
lemmings and many other species that form a major portion of the base of the food chain die off 
without the insulating effects of snow. Low temperatures that would not have been a problem 
with deeper snow became deadly. 

Expansion to the interior continued apace, and by the turn of the century competition between 
the two companies was placing even greater pressure on the resources. The Indians were caught 
in the middle of the conflict. They watched the battle and felt the effects as the land and its 
resources were hard hit. For 20 years the struggle continued, finally being resolved with 
amalgamation in 1821. 

Climate improved briefly in the first decade of the nineteenth century, but by 1809 a cooling 
trend was beginning. Much has been written about 1816, the year with no summer. It was 
originally thought that the eruption of Tambora in 1815 was the cause of this dramatic, history- 
altering year. However, climatic records show that cooling had begun in 1809, and the volcanic 
eruption occurred at the nadir of the cool period 1809-20. It is interesting to speculate on the 
impact of the volcano if global temperatures had been increasing at the time. 

Conditions in Europe had been very similar through the latter part of the eighteenth century, as 
the work of many scholars attests. Harsh conditions seriously reduced the food supply and placed 
populations under increasing stress. People responded in their traditional ways, starvation and/or 
migration. Interestingly, migration is not the choice of the majority. Lord Selkirk’s offer of 
transport to new opportunities and better conditions made the decision a little easier; but still it 
was not everyone’s choice. 

The first groups who came from Scotland under his auspices went to Prince Edward Island and 
Ontario. The best known group was the Selkirk Settlers: under the leadership of Miles Macdonell 


they arrived at York Factory in the autumn of 1811. Too late to travel south, they wintered at 
a place known as the Nelson Encampment and experienced the harshness of conditions of that 
part of the world. Things were not much better when they arrived at the Red River settlement 
at the junction of the Red and Assiniboine rivers in 1812. The cooler conditions discussed above 
had already begun, and made the early years very difficult. The harsh conditions of the "year 
with no summer" were just the beginning. In his district report for 1819 Peter Fidler writes that 
from 1816 to 1819 a severe drought affected everything and everyone: 

The spring months have sometimes storms of wind and thunder even so early 
as March within these last three years the Climate seems to be greatly changed 
the summers so backward with very little rain and even snow in winter much 
less than usual and the ground parched up that all small creeks that flowed 
with plentiful streams all summer have entirely dried up, for these several 
years loaded craft could ascend up as high as the Elbow or Carlton House but 
these last 3 summers it was necessary to convey all the goods from the Forks 
by land in Carts... 

The latter comments refer to the shift from river traffic on the Assiniboine River to the use of 
Red River Carts. This climate-induced shift is reflected in the pattern of roads and settlement 
across the prairies today. 

Consider the situation that has developed by 1817. The Selkirk Settlers have been thrust into a 
new harsh landscape. They are almost immediately confronted with severe weather that made 
things as difficult as they had been in their native Scotland. In addition, they were not welcome, 
either by the Indians, the Métis, or the fur traders. The Indians saw them as a threat to their 
traditional way of life and usurpers of their land. They were particularly concerned because they 
were suffering severely by the harsh weather and consequent lack of food. The Métis were 
already the ‘in-between’ group and had nothing to gain from the addition of another faction. 
Besides, they also benefited as a key part of the fur trade. Fur traders saw the settlers as a threat 
to their freewheeling monopolistic style. They also recognized that agriculture and fur trapping 
were potentially mutually exclusive. Amalgamation between the two companies had not occurred 
yet, and the settlers were unwitting pawns in the conflict. 

The entire situation was extremely volatile. It is not surprising that the Indians and Métis attacked 
the settlers or that the fur traders (particularly members of the North West Company) did little 
to assist them. The culmination of these conflicts was the Seven Oaks Massacre in 1817 when 
a group of Métis led by Cuthbert Grant killed 21 people. I do not think that the climate was the 
cause of this event, however, I suggest that the extreme climatic conditions and their impact on 
the economy and food supply created untenable and volatile situations. The fur trade was to 
continue for some decades as there was little useable land in northern Canada. However, the 
southern regions were irretrievably changed as the land was cleared at an increasing rate. 


Borisenkov, the Soviet climatologist, has carried out extensive research using historical sources 
to reconstruct climate over the last several centuries. He writes that, "In the climatic sense the 
Little Ice Age was highly variable both spatially and temporally. The main feature of that period 
was the frequent recurrence of climatic extremes, during which Russia suffered 350 “hungry 
years" as a result of unfavourable climatic conditions." He makes links between the climatic 
disasters, such as drought, rainy and cool summers or severe winters and the pattern of peasant 


life. The correlation between particularly prolonged harsh conditions and peasant revolts cannot 
be ignored. 

It is easy to blame historians for not considering climate as a major factor influencing important 
social events. They have not had the information - although some clues should have been 
apparent. Paintings such as those by Breughel showing winter conditions very different than today 
or Jan Griffier’s frosty painting of the River Thames with 2 feet (0.6 cm) of ice in the great frost 
of 1683, cannot be accused of artistic licence. As the picture of historic climate is reconstructed 
we should be able to reach more precise conclusions about the relationship between climate and 

The concept of climate influencing the pattern of history has suffered from the extreme distortions 
of fascism and the lack of information about actual climatic conditions. Climatic determinism is 
not the issue here, especially as it relates to human characteristics. My point is that climate affects 
the environment which has direct impact upon the food supply and economy, and therefore the 

The fur trade of North America provides an excellent opportunity to study the relationship 

between climate and the pattern of history. Records maintained by the Hudson’s Bay Company 
provide detailed evidence of the climate and its impact on the economy and lives of the people. 


The Ecology of a Famine: Northwestern Ontario in 1815-17 

Roger Suffling’ and Ron Fritz’ 


The extraordinary summer weather of 1816 has been blamed on the eruption of the Tambora 
volcano in 1815, and has been associated with various social and economic disruptions around 
the globe. In northwestern Ontario there were famines in the winters of 1815-16 and 1816-17 
among native Ojibwa people and Hudson’s Bay Company traders who relied on seven basic 
resources: moose, caribou, wildrice, potatoes, fish, wildfowl and furbearers. The first of these 
two famines was extremely severe. It resulted from an initial, non-climatically induced reduction 
in moose and caribou, and a natural cyclic crash in the snowshoe hare population. The early 
summer drought of 1815 reduced the potato crop, and late summer rain and cold ruined the 
wildrice harvest and fishing. These, combined with an early fall goose migration left both the 
Ojibwa and Hudson’s Bay Company employees starving. In contrast, the dry cold summer of 
1816 fostered production of a large potato crop, though of indifferent quality, and normal 
wildrice and fish harvests. The winter 1816-17 famine resulted primarily from deep snow 
conditions, but the Hudson’s Bay Company was able to feed many starving natives. Thus the 
cold, droughty summer of 1816, the "year with no summer", may have done as much to 
ameliorate famine among the Ojibwa people as to create it. 


Northwestern Ontario (Figure 1) is a harsh land of subarctic forests. Even now, the population 
is sparse and, in the early nineteenth century, it probably never exceeded a few thousand (Bishop 
1974). Though severe epidemics of smallpox and other diseases periodically devastated the 
Ojibwa and Cree peoples (e.g., Hearne 1791), starvation was often the most effective controlling 
factor, as the following report from Osnaburgh House (51° 90’N 90° 15’W) illustrates: 

"An Indian woman of the Crows gang came in to [sic] day with her four 
young children all much starved and with a very miserable report. Says that 
her husband starved to death two winters back since which she has been in 
wretchedness and want with four children to support. Her friends take but 
little notice of her being under the impression that she eat her husband when 
under one of the greatest of all miseries extreme starvation." (1 November 

The above is part of a daily record kept by Hudson’s Bay Company (HBC) Post Masters, and 
known as a Journal of Occurrences. It spans the years 1786 to 1911 with few interruptions. The 
journals confirm the generality of the appalling conditions described above (though this is a 
severe case indeed). Of 126 years of records, 60 include allusions to people starving (Figure 2). 

' Faculty of Environmental Studies, University of Waterloo, Waterloo, Ontario N2L 3G1, Canada. 
? Unless otherwise noted, dates quoted are from the Osnaburgh House Journals of Occurrences, originals of which 
are kept at the Hudson’s Bay Company Archives in Winnipeg, Manitoba, Canada. 









oe | 

oH ; 

Joc yp Ss IRR Se pene ed Ha | 
Neat ee { 




| : Sees Bay} 
\. ESCABACHEY i o—" Lowlandst 


te. Narie 

Figure 1: Locations mentioned in the text. 

In using the word, "starving", we intend the same meaning attached by the Post Masters: that 
somebody was unable to procure food for days at a time. Sometimes this situation would be brief 
or intermittent. On occasion it continued until death ensued. Often, however, it is apparent that 
starvation contributed to death by other means. The woman cited above struggled on in the same 
pathetic condition for two more years before succumbing to an illness: she had killed her husband 
to survive a famine. 

We have used the word "famine" to mean general starvation among people that was sufficiently 
prolonged to become life-threatening. 




: Seale 
100 | 


50 a 
os aF Sak 

oo 2 & te © 68 © &® © 8% oo ww & & oo & Oo HM & 
ce) op oe Oo @& S Ff A AV se Ww S&F yw & ® © oO EE WS 
Im |e Is co) Cs) Co} bs) (os) (o) Ce) eo} Co ea) (o} (oe) (20) Co) (00) J fes) (bo) 

Figure 2: Incidence of starvation at Osnaburgh House. The starvation index is the product of the number 
of references to starvation in the Journal of Occurrence and the maximum number of people 
recorded as starving. 

In the fall and winter of 1815-16 there was a particularly horrendous incident of famine in 
northwestern Ontario that is well illustrated by the Osnaburgh House records. It involved not only 
the native population, but also (and unusually) the better prepared HBC employees and their 
families. The 1815-16 famine was one of eight recorded that involved more than 30 people (out 
of a total population of about 200). In terms of numbers referenced in the journals, it was one 
of the worst eight incidents, and in deaths it probably ranks only second to the 1823-25 incident. 

There was another famine in the 1816-17 winter but, though it was severe, and though it may 
have involved just as many people, its consequences were not as grave as those of the 1815-16 
famine. Both of these famines are part of a prolonged series of incidents from 1810 to 1825 that 
broke the spirit and culture of the Ojibwa people (Bishop 1974). The famine series is associated 
with depletion of big game and an exceptionally cold climatic fluctuation. 

The 1815-16 famine is particularly striking because it corresponds with climatic fluctuations 

evidently caused by the eruption of the Tambora volcano in what is now Indonesia, early in 1815. 
The question that we asked ourselves, therefore, was whether the 1815-16 famine was caused by 


unusual weather conditions. To find an answer, we examined the ecology of the Ojibwa people 
around Osnaburgh House to see which climatic or other conditions normally contributed to 
famine, and to see which of these pertained immediately before or during the 1815-16 incident, 
and the lesser famine of 1816-17. 

Ojibwa Ecology and Food Sources in Early Nineteenth Century Northwestern Ontario 

Big Game 
Bishop (1974) has postulated that the Ojibwa people who lived near Osnaburgh House in the early 
1800s had moved there from around Sault Ste. Marie. 

Initially, they had been primarily big game hunters, subsisting on moose (Alces alces) and 
woodland caribou (Rangifer tarandus). Bishop believes that, after first European contact, the 
people began to make forays into northwestern Ontario in search of furbearers to use in the new 
commercial fur trade. Rival companies soon began to challenge the HBC’s hegemony over this 
and other areas. First the French and then Scots from Montreal, and American traders moved into 
the area to trade furs at their source. This induced the Ojibwa to remain on the summering 
grounds year round, but forced a radical reorganization of their hunting strategy. Originally, they 
had hunted in large groups. Now, with the need to spread out to trap beaver, they broke into 
family groups, and they used firearms to kill moose and caribou. When they were available in 
sufficient quantity, moose and caribou meat and skins were traded to the HBC, putting further 
pressure on the herds. By 1815, both species were already somewhat depleted (Figures 3, 4), and 
this was beginning to wreak hardship, not only directly in terms of food availability, but also 
because leather for mocassins and snowshoes was becoming scarce. The people’s very existence 
in this land of thinly-spread resources, was predicated on nomadic foraging, so a lack of leather 
hampered many food gathering activities, as well as in fur trapping. In a typical instance, a native 
arrived at Lac Seul asking to purchase a summer bear (Ursus americanus) skin, there being no 
caribou or moose leather: "The bear skin is for making his shoes without which he cannot leave 
his tent" (Lac Seul 28 April 1828). 

Moose (Figure 3) and caribou (Figure 4) were stalked at all times of the year, and herded into 
the water for slaughter in the summer. Deep snow slowed the animals down in winter, making 
them easier to approach, but they could easily outrun hunters on thin snow - so the latter 
condition is associated with hardship. Extremely deep snow made both hunter and hunted less 
mobile and sometimes prevented the people from traveling between various moose and caribou 
wintering grounds. 

Crusted snow gives human or other predators a marked advantage (J. Theberge 1988, personal 
communication). It must have occurred with greater frequency in the early nineteenth century as 
uncontrolled forest fires increased the proportion of open country where crusting occurs easily. 
Thus, even as the herds were reduced, the pursuit of the remaining animals may have become 
more efficient, ensuring further big-game depletion. 

Moose and caribou meat were eaten fresh, or preserved by drying in strips over a fire, or in 
pemmican (a preserved mixture of fat, berries and shredded meat pounded together). 






Figure 3: The number of moose involved in trading of meat and skins from natives to the HBC at 
Osnaburgh House. The upper line is a maximum estimate, and the lower line a minimum. 
Derivation of the data is given in Fritz (1988). 






Om OO Ow 

ea ee 


Figure 4: The number of caribou involved in trading of meat and skins from natives to the HBC at 
Osnaburgh House. The upper line is a maximum estimate, and the lower line a minimum. 
Derivation of the data is given in Fritz (1988). The maximum figure (102) for 1876 is off-scale. 



The third major food source was fish. They were hooked, speared or netted, depending on 
species and season. Several species were used including: whitefish (Coregonus clupeiformis), 
sucker (Catostomus spp.), pickerel (Stizostedion vitreum), sturgeon (Acipenser fulvescens) and 
pike (Esox lucius). At Osnaburgh House, sturgeon appear to have been particularly critical to 
human welfare. They could be readily speared and netted when spawning in the early spring - 
an otherwise lean time of year. Spring and summer were employed in catching mostly pickerel, 
pike and whitefish; and the fishery continued until the water became warm and the eating quality 
of the fish declined. Fishing resumed in the fall as water temperatures fell, when a number of 
species came to spawn in the shallows and rapids of the rivers. Fishing continued until freeze-up 
and occasionally afterwards, under the ice, but the early nineteenth century natives do not seem 
to have mastered the art of ice fishing with nets as the HBC people had. 

High water in the lakes generally signalled a failure of the fishery, especially in the fall. The high 
water could be caused by unusually heavy rainfall, cool weather, or a combination of both. Early 
freeze-up also hurt the fishery as it cut short the spawning seasons of the fish, and they withdrew 
to deeper water. 

Fish were vitally important at northwestern Ontario HBC posts during winter - especially 
whitefish and, as soon as the weather became cold enough to store fish, they were netted 
intensively. In times of native starvation these fish were distributed to Ojibwa begging at the 
posts, as long as the HBC’s own supply of stored or fresh fish remained assured. The motivation 
was partly charitable but hinged too on the economic need to preserve the lives and health of the 
beaver trappers who were the lifeblood of the Company’s activities in these parts. 


The fourth major native food resource was wildfowl -- primarily geese. Both Canada Geese 
(Branta canadensis) and Snow Geese (Chen caerulescens) were shot, as well as a variety of 
ducks. At Osnaburgh House, wildfowl first appeared in April, moving north to the Hudson Bay 
Lowlands and beyond on the turbulent edge of the retreating Arctic air mass (Ball 1983). If snow 
lies on the coastal marshes of Hudson Bay at the time when goose eggs should be laid, a lower 
proportion of females than normal actually lays eggs. In addition, average clutch-size is reduced. 
Thus the cohort of young geese produced is small, as happened in 1967. Fall migrants then prove 
relatively sparse, as do those birds returning the following spring. If fall came early, sending the 
birds south too soon for native needs, then people had a longer time to wait between fall and 
spring migrations. The people were often starving in late winter, so that the return of the geese 
was awaited with eagerness by both natives and HBC men. 


Wildrice was the only staple vegetable of the largely carnivorous Ojibwa. (The same cannot be 
said of the HBC men who also grew potatoes and some lesser crops). Wildrice is an annual 
aquatic grass found in slow-flowing rivers and shallow lakes (Dore 1969, Suffling and Schreiner 
1979). It sets seed in late summer and is harvested in late August or early September in 
northwestern Ontario. The seeds, which were fermented, hulled, dried and stored for winter use, 
were a good hedge against starvation. High water in mid- to late-summer - especially rising high 
water - is disastrous to the crop. Also, windstorms can scatter the grain before it is harvested. 



Snowshoe hares (Lepus americanus), usually called rabbits in the journals, were also an important 
food item. They never appear to have been a preferred food (Bishop 1974), but were snared in 
hard times when other victuals were lacking. Hare pelts that were the by-product of this activity 
were traded, but only commanded a minimal price at the HBC posts. Alternate freezing and 
thawing in winter made rabbit snaring impossible (Lac Seul, 3 February 1825). 

The snowshoe hare exhibits a remarkable eight- to nine-year cycle of population density 
(MacLulich 1937, Elton and Nicholson 1942). It is notable that most of the peaks in human 
starvation at Osnaburgh House appear in the year after the crash of the hare population 
(Figure 5), a relationship which is statistically significant (X2, P<0.01). Thus snowshoe hare 
scarcity could precipitate famine. 


Furbearers were the means by which the Ojibwa obtained non-local commodities such as iron 
knives, hatchets, guns, blankets and rum. The species trapped or hunted include marten (Martes 
americana), otter (Lutra canadensis), fisher (Martes pennanti) and lynx (Lynx canadensis), but 
the most important was beaver (Castor canadensis). Where they were available in large numbers, 
as at Lac Seul, muskrats (Ondatra zibethicus) were also very important in total, and as beavers 
were depleted, muskrats assumed an increased economic significance. Although beavers and 
muskrats had the added advantage that the carcasses were edible, generally furbearers did not 
contribute greatly to human nutrition. Their significance in this context is in how they influenced 
the pattern of trapping and hunting of other animals. 


As atule, natives in early nineteenth century northwestern Ontario did not grow potatoes, though 
there were a few individual attempts. This vegetable was, however, a staple of the HBC posts. 
Potatoes were planted in early May at Osnaburgh, and harvested in mid- to late-October. 

The potato crop was highly variable. It appears to have suffered after hot, dry summers, and was 
of low quality if an early fall frost damaged the tubers. Though spring frosts were damaging to 
the top growth and may have reduced the yield, they do not seem to have been as serious a 
problem. Every few years, there were also damaging epidemics of "grubs" (so far unidentified). 
As with fish, potatoes were given to starving Ojibwa coming to the post for assistance - at least 
for as long as there was no threat of starvation to HBC employees themselves. 

The Annual Cycle of Ojibwa Subsistence 

The annual cycle of Ojibwa subsistence in the early nineteenth century (Figure 5) is not totally 
dissimilar to the modern pattern described by Sieciechowicz (1977). The Ojibwa year can be 
thought of as beginning in late September to October when natives arrived at the HBC post to 
obtain their outfit for the coming winter. At this time, goods were normally obtained on credit, 
a process described as "taking debt" or "outfitting". The fall fishing and goose harvest more or 
less coincided with taking debt. Then, as the weather hardened, and the furbearers came into 
prime pelage, there was a concentrated effort to trap - and especially after the first snows. The 
amorous bull moose could be readily killed at this time as they could be called in by a hunter 
using a birch bark trumpet to initiate a rival or a cow moose. 


When the large lakes froze, usually in early November around Osnaburgh, winter began in 
earnest. By now only big game, hares and occasional grouse (Canachites canadensis, Bonasa 
umbellus, and possibly Pedioecetes phasianellus) were available. Since stored fish and wildrice 
gave out, often about the beginning of January, people had to since rely entirely on meat and fat. 
If starvation arose, it became apparent in the journal entries at this time, and it would become 
more severe after the onset of really bitter weather. The starving time of winter could be 

alleviated or avoided if large game abounded, if hares were present in large numbers, or if spring 
came early. The converse was also true. 


way JULY 

APR . 
\ AUG 

| | \ 
ag \ | 


Aelia Pica eae f dut-; 

ih Fitting 
FEB, on ae dae 
JAN” Fur-Bearers—— NOU 

Figure 5: The annual cycle of early nineteenth century Ojibwa people living around Osnaburgh House. 


If people were not starving severely, there was a second burst of fur-trapping activity in the 
relatively mild weather of late winter and early spring. 

Normally in mid-April the first geese arrived, and sometimes set off what can only be described 
as a hunting frenzy at the HBC posts! At Lac Seul in 1828, for instance, the Post Master gave 
all his people three days off to hunt geese,"... in the hopes of their setting to work afterwards". 
Geese not only provided relief from starvation, but for the HBC men in particular, brought a 
welcome rest from the six-month monotony of potatoes and whitefish. 

In late winter Ojibwa appeared at the posts to redeem their debts, to socialize, and to drink. It 
was a time when the HBC Post Master anxiously awaited the fur harvest, and when the extent 
of any lethal starvation became apparent through the non-arrival of families from the forest. 

With the break-up of ice on the lakes and freshets of meltwater in the rivers came the spearing 
of sturgeon and pickerel as they spawned at the base of rapids. Bears (Ursus americanus) too 
came for the fish, and could be readily trapped then, if they had not been found in their 
hibernation dens and speared. They were generally too lean in the spring to provide much meat 
or fat. Fishing continued throughout the spring, supplemented by game hunting, as well as by 
collecting a variety of fruits and berries, and possibly birds’ eggs. 

In August the water became too warm for profitable fishing, but at the end of the month the 
wildrice harvest began. This was another time that brought people together, and it was closely 
followed by fall fishing and acquiring new outfits at the Posts. 

The pattern described above is typical, but each year presented a slightly different situation, and 
the resources available around each trading post differed slightly. Osnaburgh had more sturgeon, 
Lac Seul had more muskrats and wildrice, etc. The trading policies of the HBC and its rivals, 
the weather, as well as availability of food and furs all varied enormously over time, and have 
been discussed at length by Bishop (1974). 

Factors that Precipitated Starvation 

The factors causing or excacerbating starvation are summarized in Table 1. Severe starvation 
might be avoided if only one or two factors were unfavourable in a given year, but if several 
coincided, then people would suffer accordingly. Most of the factors have been discussed above, 
but the incapacitation of hunters needs comment. 

Injury, sickness, or death of menfolk was a constant peril to family groups - it could deny them 
access to big game. If the women and children’s occupation of snaring hare was unavailable 
because of a hare population crash, then starvation was bound to follow. Repeated freeze-thaw 
cycles that prevented hare snaring sometimes had the same effect. 

The 1815-16 Famine 
January and February 1815 were fairly typical for the time of year - dry and cold. This weather 

persisted into April which, coupled with north winds, kept most of the geese and ducks from 
arriving. A few came, however, on 15 April - about the usual time. 


May, too, proved very cold at first so that the Post Master remarked that the weather had "more 
the appearance of March than of May" (8 May 1815). The main body of geese arrived only on 
14 May, a month late. The keeper of the journal considered that the latter half of May was warm 
for the time of year, and the lake ice broke only a little late. 

Early June was judged to be warm for the season, but the latter half was rainy and cold. This 
weather persisted into early July until it became very hot during 9-13 June, and then again from 
24 July to 10 August. The warm spell was sufficiently drying that the writer states: 

“Hookamarshish informs me that all his furs were burned, he says that he was 
going to move to another place and he forgot to put out his fire and so it set 
fire to the woods and burned all his furs." (15 August 1815). 

Table 1: Factors Contributing to Starvation among the Early Nineteenth Century Ojibwa around 
Osnaburgh House. 

Causal Factor Resource Affected 
Climatic Contributors 

High water in summer Wildrice, fish 

‘Rising water in summer Wildrice 

Thin snow Moose, caribou 

Very deep snow Moose, caribou 

Cold, damp spring on Hudson Bay Geese 

Droughty summer Potatoes 

Freeze/thaw in winter Snowshoe hare 

Non-Climatic Contributors 

Increased forest fires due to Moose, caribou 
fur trade 

Overhunting due to fur trade Moose, caribou 
Cyclic population crash Snowshoe hare 
Injury, sickness or death Moose, caribou 
of menfolk 


On 22 August came a sudden cooling with rainy, stormy weather accompanied by NW and E 
winds. These conditions persisted unabated until 17 September. By late September it was apparent 
that both the fall fishing and the wildrice harvest had failed on account of high water in the lakes 
and rivers: 

"The water being so remarkably high at this place the Indians is not made 
any rice worth while so that I have only got 64 gal in all. So that I am much 
afraid of starving in the winter as there is no fish to be got here when the 
water is high in the fall. Am sorry to inform you that this is a very poor 
place for most everything. There is no beaver nor moose to the indians to 
hunt and most of them were starving when I seed them but are all off now 
to hunt." (Letter from James Slatter at Escabachewan 23 September 1815 to 
the Master of Osnaburgh House). 

At Osnaburgh, the problems were compounded by a lack of fishing twine and of available labour. 
Such fall starvation was unusual, but the people were probably cheered by the early arrival of 
the bulk of the fall geese on 1 October. In reality, this worsened matters for, with the early 
passage of wildfowl, the impending winter starvation was to last longer. 

The potato harvest at Onsaburgh House was 77 kegs, down 20 from the previous year, so that 
the HBC people entered the winter with very little food to spare for visiting Ojibwa. 

The first snow came on 22 October and the lake froze on 7 November, a trifle early. The 
subsequent ice fishery failed as miserably as had the fall netting. The first half of December was 
very cold and the the latter half mild. 

Starvation is first mentioned again in the journals in December, and by late January 1816 it was 
general among the natives, even appearing at the HBC fishing outposts: "The men are already 
feeling the iron hand of want." (Osnaburgh House, 30 January 1816). 

In February, which was cold even for that time of year, natives arrived at the Post both frozen 
and starved; but others coming from the north were heavily laden with furs and apparently well 
fed (according to Bishop (1974) there were still moose to be had in that quarter). 

By 23 March, the potato ration for HBC people had been cut to two gallons per week (instead 
of the usual three three gallons), and by 25 March everybody was sent out to hunt or fish since 
the daily ration was, by that time, one small fish. The men at the marsh outpost of Osnaburgh 
were now too weak even to go to the House for food and one fellow, reduced to eating fish offal, 
became very sick. 

Very few natives had visited the Osnaburgh House during the winter, either because they were 
too weak to travel or possibly because word was out that there was no food to be had there. 

April was very cold, and the digging of the potato garden at Osnaburgh House began two weeks 
late (on 30 April) as a direct consequence. The famine finally broke with the arrival (three weeks 
late) of the first geese on 4 May. Sturgeon did not begin to spawn until 29 May - two weeks later 
than usual. 


June 1816, likewise, was very cold with a hard frost on 4-6 June and another on the 23-27 June. 
On both occasions the gardens were badly frosted. On the latter, there was one-quarter inch of 
ice in the bottom of the canoes pulled up on the shore. This must have been a dry month as the 
lake fell six inches in three weeks. 

July was cool and rainy with mostly NW winds, and this weather continued into the first half of 
August. In spite of this, the water remained low in the Albany River, suggesting perhaps that 
there had been little snow in the previous winter, and that water in the marshes must have been 
low all summer. 

On 18 August, there was snow - an unheard-of event in this month, and with continuing cold . 
weather the geese were already flying thick by 15 September - a month early. During 25-30 
September there was a gale, remarkable not so much for its ferocity but for its duration. Its winds 
tracked from E to S, to SW. In contrast with 1815, there is no indication that the wildrice harvest 
or the fall fishing were other than normal at Osnaburgh House. 

The ground froze by the 3 October (about three weeks early), so the HBC people were caught 
unprepared and the potatoes were frozen in the ground. In spite of this, they harvested 190 kegs, 
a large crop, though evidently of indifferent quality on account of the frost. The rye plants were 
six-feet tall but the grain was still green, and never had a chance to ripen. The whole crop was 

The lake froze a little early on 9 November and the weather continued cold until the second half 
of December which proved mild and snowy. 

January 1817, and the rest of the winter, were cold with heavy snowfall, which was in marked 
contrast with the previous cold, and apparently dry winter of 1815-16. 

There are almost as many citations of starvation in the 1816-17 journal as there had been for 
1815-16, but there is little indication of the grinding life-threatening severity of famine which had 
overtaken people in the previous year. Apparently the 1816-17 starvation touched only the native 
people, and many of them were visiting the post for handouts. 

Discussion and Conclusions 

It would be easy to rush to the conclusion that the Tambora eruption of 1815 explains any 
unusual weather patterns in the subsequent couple of years. As the discussions of the "Year 
without a summer? Climate in 1816" conference demonstrated, it is difficult to unequivocally 
establish causal connections, even though there is a suspicious conjunction of climatic dislocations 
around the globe. At Osnaburgh House, the unusual conditions were: the sudden cool, wet end 
to the hot, droughty summer; the long, dry, cold winter of 1815-16; and the dry cold summer 
of 1816. 

Likewise, the mere existence of famine in 1815-16 is not proof, per se, of the human ecological 
consequences of the Tambora eruption, or even of the effects of the harsh weather of these years. 
In reality, several factors contributed to the famine at Osnaburgh House and elsewhere in 
northern Ontario. They include the following ecological factors: 


1. Depletion of moose and caribou herds by overhunting (Bishop 1974), or possibly by habitat 
change through forest fires in the late 1700s. Both of these are associated with the expansion 
of the European fur trade. 

2. A cyclic crash in the hare population between spring 1814 and spring 1815 fur returns. 

Neither of these two factors are climatic, but they are primary causes in the sense that they set 
the stage for the other events. The fate of many natives was sealed by other phenomena that were 
indeed climatic, namely: 

1. Failure of the wildrice harvest in 1815 due to high, rising water levels in late summer. 

2. Failure of the 1815 fall fishery due to high water, and failure of the ice fishery, possibly for 
the same reason. 

3. The small 1815 potato harvest resulting from the early summer drought and hot weather. (The 
potatoes were grown in a dry, "hungry", sandy soil that warmed quickly in spring but was 
very vulnerable to drying.) Thus the HBC had few or no potatoes to spare for the natives 
during the famine. 

4. The early-fall goose migration in 1815, and the late-spring migration of 1816. The former 
ensured that the natives entered the winter in poor nutritional condition, and the latter 
prolonged their suffering in spring. 

One last factor was the lack of HBC labour for fishing, and lack of fishing twine in the fall 1815. 
These are minor economic or logistic causes. 

The summer of 1816 was in marked contrast to that of 1815. Though there were late spring frosts 
and summer snow, the cool weather actually appears to have helped the potato crop. Likewise 
the droughty conditions kept water levels low and assured at least a normal fish and wildrice 
harvest. Spring frosts had few harmful effects, and the fall frosts did some damage to potatoes, 
but not enough to be serious. 

It is as impossible to say that climate alone caused the 1815-16 famine as to claim successfully 
that non-climatic factors were responsible. People relied on seven major resources. Moose and 
caribou had already been depleted by 1810, and the worsening climate of 1810 to 1817 was the 
trigger for a series of famines that were only alleviated by the hare "high" of 1814. The 
subsequent crash of hares reduced the resources available to four: wildfowl, wildrice, fish and 
potatoes. The high water of 1815 knocked out two of these - fish and wildrice - leaving only 
potatoes and wildfowl - and even the potatoes were reduced by early drought. A catastrophe then 
became inevitable. 

The fall of 1816 was different. Big game and hares were still scarce, but the water remained low, 
ensuring a wildrice and fish harvest. The cool summer evidently favoured the HBC with a large 
potato crop, although possibly more had been planted as a reaction to the previous famine. On 
the other hand, the October frost impaired the quality of the crop. The two frosts of June had 
evidently had little effect, although one might easily jump to the conclusion that they had caused 
the smaller famine of 1816-17. In reality the potatoes were the saving grace for a native 
population that probably was still reeling from the physical and psychological impact of the 
1815-16 famine. The winter or 1816-17 was a year of starvation, but was not as serious as its 


predecessor. If anything, it must have been deep snow that limited travel to new hunting grounds 
or to the HBC for charity, that caused most hardship. 

We conclude that the extraordinary cold, dry weather in the summer of 1816 may actually have 
done more to prevent famine than to create it. The cold, wet end to the summer of 1815 was, 
however, the proximal cause, but only the proximal cause, of the 1815-16 suffering. 

The postscript of the famine is as interesting as the event itself. Between 1819 and 1820, the 
number of both moose and caribou traded at Osnaburgh House rose dramatically (Figures 3, 4). 
Perhaps enough hunters perished that the predation pressure on the herds was reduced and they 
started to increase. Deep snow in 1816-17 may have reduced hunting pressure with the same 
effect. Whatever the causes, moose and caribou then persisted until the mid-1820s before 
succumbing to hunting or other pressures. Thus one effect of the 1815-16 incident had been to 
prolong the survival of big game that were so important to the people, even as it helped to 
destroy their will and culture. 


The research for our paper was conducted while one of us (R.F.) held a Natural Sciences and 
Engineering Research Council Undergraduate Internship. We thank John Theberge and 
Harold Lumsden for their advice concerning ungulate and goose ecology. 


Ball, T. 1983. The migration of geese as an indicator of climate change in the southern Hudson 
Bay region between 1715 and 1851. Climatic Change 5:85-93. 

Bishop, C.A. 1974. The Northern Ojibwa and the Fur Trade, An Historical and Ecological Study. 
Cultures and Communities Series. S.M. Weaver (general ed.). Holt, Rinehart and Winston, 

Dore, W.G. 1969. Wildrice. Canada Department of Agriculture Research Branch, Plant Research 
Institute Publication 1393. Ottawa. 84 pp. 

Elton, C.S. and A.J. Nicholson. 1942. The ten year cycle in the numbers of the lynx in Canada. 
Journal of Animal Ecology. 1:215-244. 

Fritz, R. 1988. Moose and caribou population decline in N.W. Ontario boreal forests of the 
Osnaburgh House (HBC) trade area: 1786-1911. Senior Honours Essay. Department of 
Geography, University of Waterloo, Ontario. 

Hearne, S. 1791. A journal of observations made on the journey inland from Prince of Wales Fort 
in latitude 58°50’ North to latitude 72°00’ Beginning 7th Decr. 1770, ending June 30th, 
1772 by Samuel Hearne. Manuscript in British Museum Library, London, U.K. 

MacLulich, D.A. 1937. Fluctuations in the numbers of the varying hare (Lepus americanus). 
University of Toronto Studies, Biology Series 43:1-136. 


Sieciechowicz, K. 1977. People and land are one: an introduction to the way of life north of 50°. 
Bulletin of the Canadian Association in Support of Native Peoples 18(2):16-20. 

Suffling, R. and C. Schreiner. 1979. A Bibliography of Wildrice (Zizania species) Including 

Biological, Anthropological and Socio-economic Aspects. University of Waterloo School 
of Urban and Regional Planning, Working Paper 5, Waterloo, Ontario. 

Paily | 

The Development and Testing of a Methodology for Extracting Sea-Ice 
Data from Ships’ Log-Books 

Marcia Faurer! 


The severity of the weather in 1816 in the Hudson Strait and Hudson Bay regions became 
apparent through the reconstruction of sea-ice conditions for the period 1751-1870. The sea-ice - 
data were derived from an exceptionally large collection of ships’ log-books. Current research 
is focusing on the development of a reliable methodology for use in further environmental 
reconstructions covering this period using historical documents. 

This research applies a methodology called content analysis which was developed by the Social 
Sciences for extracting meanings from human communications in an objective manner. This 
technique has the ability to reduce the subjectivity inherent in the interpretation of historical 
documents by testing the level of reliability of the procedure that is used to extract sea-ice data 
from log-book descriptions. In this study, tests have been applied throughout the development of 
the methodology with the goal of devising an objective procedure. These tests also reveal the 
degree of detail that a particular source can reliably provide, as well as helping to reduce the 
difficulties associated with calibrating the historical terminology against the contemporary sea-ice 


Although content analysis (CA) has been used widely in the application of historical documents 
as proxy sources for climatic reconstructions, this methodology has not been applied to its fullest 
potential. This is primarily due to the general omission of its strongest attribute, which is the 
ability to test the reliability of the methodology that is used. This aspect of CA is not merely an 
option that may or may not be applied, it is actually an integral part of the CA process. Without 
this means of evaluation, the interpretation of historical texts may be guided by predetermined 
decisions about the information required for the reconstruction instead of by the information that 
the documents can objectively provide. 

This case study was conducted to test the applicability of an objective methodology for extraction 
of environmental data from historical documents. The format of CA was followed closely by the 
repeated application of reliability tests. Whereas this reduced the information obtained, it insured 
that the derived data were obtained to a measured and acceptable level of reliability. 

Data Sources and Background Information 
The eighteenth and nineteenth century log-books of the Hudson’s Bay Company are a potential 
source for a wide variety of environmental data. Although temperature readings were entered in 

the log-books, they appeared sporadically throughout the period of record (1751 to 1870). Wind 
directions were given on a fairly regular basis as well as other meteorological phenomena, 

' Department of Geography, University of Manitoba, Winnipeg, Manitoba R3T 2N2, Canada. 


however sea ice was chosen to be the focus of this study. While sea ice is not a meteorological 
element per se, it is a visible expression of several environmental factors. It was also selected 
because it posed a clear and present danger to the success of the voyage and to the lives of the 
crew. Therefore, it was anticipated that any event or observation related to sea ice would be 
faithfully recorded in the log-books. 

Sea ice was encountered by the ships in Hudson Strait and Hudson Bay during the westward 
portion of a voyage between England and the Hudson’s Bay Company’s bayside posts. These 
locations and the routes of the ships are shown in Figure 1. Each year, the Company dispatched 
a small convoy of ships to supply these remote posts and to bring trade goods back to England. 
This collection of log-books provides an unbroken record of sea-ice descriptions that can be 
cross-checked because each of the ships in the convoy kept at least one log-book. Figure 2 is a 
reproduction of a log-book page showing the meticulous way in which the environmental 
observations were recorded. Fortunately, this format and the vocabulary used in the log-books 
remained virtually unchanged throughout the entire period of record. 

To wes ee 


Factory HUDSON 

Figure 1: Sea currents and Hudson’s Bay Company sailing route. 


/ Lisesstes Wade Ualbor Boer ahs Praarday of Lb ae wes 
aL LL Wreath tare hot fac 
lal lee © leet aol 
att] lm] a |e 
2 eee ee ee 
At | leno 7 |e 
cf ha tee | 
71 a Ea 2 
AL |g er en 
ih ee eel les = 

Tiff ome [Ol LL] Ae Foy 7 7 aga 
aL ee em Posy lay oH 
AMA eT 2 Yueh aa] fo halt foc 3rmPe 

| Alla Aes ny he Bak l 

att || iaealsent : 
AL) Ws wen nl Tr Ty ng 
A et teal oh foe ME 
AL = eee Za 
OILS eR VED Tan f Bnd Lan Mand 

fol fp Bos et) ee A hirer 

Figure 2: Sample log-book page. 


Problems of Interpretation 

The problems arising from this type of data source are twofold. The first aspect of the problem 
lies in the interpretation of the descriptive accounts that used a vocabulary which was different 
from the current terminology. Secondly there is the difficulty that arises in conversion of 
qualitative or verbal descriptions into quantitative or numerical data that can be compared with 
contemporary records. When properly applied, CA serves to resolve these problems reliably. 

Although these historical sea-ice descriptions used the same terminology throughout the period 
of record, their conversion into a sea-ice index was complicated by two factors. The first is that 
the Hudson’s Bay Company did not provide a dictionary of the terms since they were not 
originated by the Company but were passed down through two generations of ships’ captains. 
Therefore it is not possible to translate the terms directly from an historical lexicon to their 
current counterparts. Secondly, even though the individual words may have been consistent 
through time, they were not used in isolation, but rather they were used in phrases and sentences. 
As a result, their meanings changed in relation to the context in which they were used. An initial 
survey of the log-books resulted in a compilation of 90 representative phrases, an example of 
which is given in Table 1. 

Consequently, the system of classifying these phrases had to allow for any number of 
combinations of terms. In cases where there is a finite set of terms or phrases, it is possible to 
establish specific rules and guidelines for their classification. In this case however, this approach 
would not have been sufficiently flexible. As a result, it was obvious that a certain degree of 
subjectivity would be an unavoidable component of the analysis. The goal therefore, was to 
devise an objective system of categories. 

Figure 3 illustrates the CA procedure in the role of a translator from the written descriptions into 

numerical ice data. Even though the specifics of CA are usually designed to suit the needs of each 
individual research project, there is an established general format (Figure 4). 

Content Analysis 

Qualitative Description Derived Quantitative Data 

Figure 3: Conversion of qualitative data to quantitative data through content analysis. 





Sea Ice Processes 
in Hudson Bay 











Derived Sea Ice Data 


Figure 4: Content analysis plan. 

Dep ip4 

Table 1: Representative Eighteenth and Nineteenth Century Sea-Ice Phrases. 

1. Ice open and heavy 

2. Ice close but much smaller 

3. Sailing among heavy straggling ice 
4. Pieces of ice 

5. Passing thro’ a deal of sailing ice 
6. Passing thro’ close ice 

7. Heavy close packed ice 

8. Saw some ice 

9. Fast beset among close small ice can’t move 

The formulation of the research question is the first of a series of decisions that are made 
throughout the procedure. The crucial nature of this decision is due to the fact that CA is a linear 
process in which each step is the logical outcome of the previous step. Should an error be made 
anywhere along the line, then it will be carried throughout the entire analysis and may even be 
intensified in the process. The research question is based on two sources of information: the 
communications content and a body of theory. In the case of this study, the ships’ log-books of 
the Hudson’s Bay Company provided the communications content, and the Hudson Bay sea-ice 
processes provided the background theory. 

The set of procedures that follow from the identification of the question (Figure 4) comprises the 
core of this research. The steps followed here are repeated until an acceptable degree of reliability 
has been attained. This means that a group of independent researchers who apply the same 
methodology using the same data will consistently produce the same results. Basically, this is a 
process of redefining categories with the goal of reducing the level of subjectivity. Once this has 
been accomplished using a sample of the data, the categories can be applied to the entire body 
of descriptions, and the reconstruction can be made and tested for validity. 

This research followed the plan presented in Figure 5, and was divided into three phases. The 
goal was to objectively develop a set of categories into which the log-book descriptions could be 

Phase I 
The first phase involved 56 randomly selected log-book pages that included samples from the 
120-year period, and five coders who all had considerable experience with CA and the log-books. 

This phase was essentially experimental. A set of five categories and codes (Table 2) was 
intuitively derived, and was based on the maximum amount of sea-ice information that was 


Figure 5: Research plan. 


- 5 Categories : Intuitively derived code 
- 3 Coding Units : Day, Entry, Word 

- No Coder Training 



- 5 Categories : Code derived from 
Observers’ Manual 

- 1 Coding Unit : Watch 
- Definitions & Diagrams Provided 

T 0 



- 4 Categories : Code derived from 
Phase II 
- 1 Coding Unit : Watch 

- Definitions, Diagrams, Photos 



Tests Confirm 

desired for the reconstruction. Each of the five major headings required the coders to make a 
dichotomous decision about the meaning of the log-book transcription. Besides the 10 codes, the 
coders were given the option of indicating those transcriptions that did not provide enough 
information to make a decision regarding the particular category. This was an important aspect 
of the code insuring that each decision made by the coder was done with some degree of certainty 
and was not a forced response. It also allowed decisions to be made later about the type of 
information that could be obtained from the log-book descriptions. 

The first phase of coding was actually divided into three sections based on the unit of transcribed 
information that was coded (coding unit). In the first section, the coders were required to provide 
a five-digit code (one from each of the five categories - see Table 2) for each of the 56 days. In 
this way, a day was treated as one block of information or coding unit. This process was repeated 
by each coder five times so that their level of consistency could be determined. The second 
section involved the coding of each individual hourly entry for the same 56-day sample (a total 
of 261 entries), and again this process was repeated five times. The third coding unit was 
comprised of a list of 81 individual words, 24 of which were direct descriptions of sea ice, and 
57 described the navigational activities employed to deal with the ice (e.g., grappling, tacking, 
rounding). These words were only coded twice. 

To evaluate the reliability of the code, the coding units, and each coder, percentage agreements 
were calculated. This is the most common and rudimentary method of calculating reliability. It 
is a considerably less-than-ideal approach because it is biased in terms of the number of 
categories and coders, in such a way that the fewer of each the higher the percentage agreement 
is likely to be. The highest average intercoder agreement (among the coders) was 53.8% which 
was achieved when the day coding unit was used. The highest intracoder agreement (consistency 
level for each coder) was 80% for the entry coding unit, although all three showed high levels 
of consistency (day=68% word=70%). It is important to note however, that a large proportion 
of the agreements was due to agreements that there was not enough information. 

Phase II 

The second phase was derived only slightly from the findings of the previous phase. Again there 
were no compulsory categories so that a coder could judge that the information was insufficient 
to make a coding decision. As a result of this option in the first phase, it was possible to conclude 
that sea-ice concentration was the only category in which a decision was possible as much as 70% 
of the time. 

Phase II introduced a new coding unit: the seamans’ watch, which was more in context with the 
log-book format since the entries were actually summaries of the six four-hour watches listed in 
Table 3. Therefore the coders were required to provide a code for each watch per day (whether 
there was an entry or not). Another difference between the two phases was that the second code 
was not intuitively derived nor did it evolve from the first coding system. Instead, it was based 
on the terms and definitions found in The Ice Observer’s Training Manual (Environment Canada 
1984). Since the final goal was to create a sea-ice reconstruction that could be compared with 
current records, the logical approach was to use modern definitions in developing the categories 
and codes, and in the coding process itself. The second set of categories and codes is given in 
Table 4, and with this the coders were also given definitions and diagrams (from the Observer’s 


Table 2: Phase I Code.! 



Ice not present in vicinity of ship 
Ice present in vicinity of ship 


2 = Small area covered by ice (<50%) 
3 = Large area covered by ice (>50%) 


4 = Ice cover highly fragmented 
5 = Ice cover not highly fragmented 



Thin layer of ice 
Thick layer of ice 



8 = Ice in motion 
9 = Ice not in motion 

' (No compulsory codes). Coding units: Day (5x), Entry (5x), Word (2x). 

Table 3: Phase II.! 

Watch Time 

1. Afternoon Noon - 4:00 p.m. 

2. Dog 4:00 p.m. - 8:00 p.m. 
3. First 8:00 p.m. - Midnight 
4. Middle Midnight - 4:00 a.m. 

5. Morning 4:00 a.m. - 8:00 a.m. 

6. Forenoon 8:00 a.m. - Noon 

' Coding unit: seaman’s watch (3x). 


Table 4: Phase II Code.! 

A. Concentration 
B.  Floe Size 

C. Openings 

D. Arrangement 
E. Motion 


OS ie 

CV Si ids ae 

eee ae 

Ice Free 

Open Water 
Very Open Ice 
Open Ice 

Close Ice 
Very Close Ice 
Consolidated/Compact Ice 

Giant Floe 
Vast Floe 

Big Floe 
Medium Floe 
Small Floe 

Ice Cake 
Small Ice Cake 


Open Lead 
Blind Lead 
Shore Lead 
Flaw Lead 

Ice Field 




Ice Edge (compacted) 
Ice Edge (diffuse) 
Concentration Boundary 


. Shearing 

' No compulsory codes. Coding units: seaman’s watch (3x). 


Manual) for the terms. Each coder applied this system to all of the watches on three separate 
occasions so that their consistency could be determined. 

The analysis of these sessions followed a more complex process that eliminated biases inherent 
in the use of percent agreements. This was appropriate here because the development and 
application of the categories was more structured than in Phase I. In this case, Krippendorff’s 
agreement coefficient was calculated by using the following equations. 

Q= 1 o Dy 
Where: a = agreement coefficient 
Do = observed disagreements 

De =expected disagreements 

b c SS, 
Where: Xpe = number of disagreements in a matrix of 

category codes (or coders) 

x, = total of the marginal entries 

> ae? ,§ 
LL ee me 

Where: xp x. = the products of all possible marginal 
entries of the matrix 

x =the marginal total 

m  =number of category codes (or coders) 

The resulting coefficient is a number between 0 and 1 which, when multiplied by 100, gives the 
percentage by which the agreements are better than chance. Therefore, when the coefficient is 
0, then any agreement is completely by chance. When the coefficient is 1, the agreements are 
based entirely on the coders’ judgements with no degree of chance. 

When this was applied to the intercoder agreements (for category A - Concentration), the 
coefficient was 0.468 or 47% better than chance. The average intracoder agreement was 0.591. 
It was then decided that these figures could be improved by modifying the categories since the 
problem was not due to unskilled coders. One of the many advantages of this agreement 
coefficient is that it can also be used as a diagnostic device in the restructuring of the categories 
(or reselection of coders, if necessary). 


When the cause of low coefficient values is due to a problem with the categories, it is usually 
because the distinctions between the codes are not sufficient. This can be remedied by combining 
those codes that are most frequently confused with one another. Figures 6a-d provide an example 
of this testing procedure. Figure 6a is the basic matrix for the seven codes in the concentration 
category. The numbers in the cells indicate the frequencies with which each was used, so that all 
of the diagonal entries are the numbers of agreements among all five coders for each code, and 
the off-diagonals are the disagreements. The coefficient for this matrix was calculated to be 
0.468. Figure 6b shows the matrix if it was collapsed into two codes: no ice (1) and ice (2-7). 
Intuitively, this would be expected to substantially increase the agreement coefficient. However, 
because there is no longer a bias in favour of fewer categories the value was increased by only 
0.086. Figure 6c shows another regrouping into three codes: no ice (1), general ice descriptions 
(2-6), and complete ice coverage (7). This raised the coefficient by only 0.016. Finally, the codes 
were regrouped (Figure 6d) into four codes: no ice (1), open ice (2-3), close ice (4-5), and 
consolidated ice (6-7) and this increased the coefficient by 0.203 so the value became 0.689 
(almost 70% better than chance). It should be stressed here that the coefficients in Figures 6b, 
c, and d were all calculated from the original matrix and not by recoding. This process was also 
applied to categories B (floe size) and D (arrangement) with increases in the coefficient of 0.333 
and 0.479 respectively. 

Phase III 

The coding system for this phase resulted directly from the regroupings discussed above and is 
presented in Table 5. Category C (openings) was omitted due to infrequent usage by the coders, 
and the other four categories were regrouped as illustrated by comparing Tables 4 and 5. This 
coding session was only repeated twice because the coders’ consistencies had been sufficiently 
tested by this point. The same definitions and diagrams were used here as in Phase II, the major 
difference being that category A (concentration) was compulsory. That is, a code designation was 
required for this category for every watch of every day. The agreements were analyzed as in 
Phase II. The coefficients for category A are given in Table 6. Because the other three categories 
were so rarely used, agreement coefficients were calculated only for category A. 

Two observations are clear from Table 6. First, regrouping raised the agreement level by 21%. 
Secondly, although the averages of the coefficients for Phase III were lower than for the Phase 
II regrouped figures, they differed by only 2%. Therefore it is possible to use the calculated 
regroupings as a prediction for the Phase III coding agreements, and the third phase of coding 
could actually be eliminated. 

It was concluded that the Phase III coding system produced an acceptable level of reliability since 
an average of only 37% of the agreements were made by chance and the remaining 63% were 
reliable agreements. As a result, the seaman’s watch and the concentration category were adopted 
as the basis for the sea-ice reconstruction. 



Agreement Coefficient = .554 
Categories : No Ice & Ice 


SY ee 

1 | 



Agreement Coefficient = .484 Agreement Coefficient = .689 

Categories : No Ice, General Ice, Consolidated Ice Categories : No Ice, Open Ice, Close Ice, Consolidated 


Figure 6: Coincidence matrices and agreement coefficients. (a) Basic matrix; (b) Two-category matrix; 
(c) Three-category matrix; (d) Four-category matrix. 


Table 5: Phase III Code.! 
A. Concentration 

Ice Free 

Open Water/Very Open Ice 

Open Ice/Close Ice 

Very Close/Consolidated/Compact Ice 


B. Floe Size 

Small Ice Cake 

Ice Cake 
Medium/Small Floe 
Big Floe 


C. Arrangement 

Strip/Diffuse Ice Edge/Concentration Boundary 


Ice Field/Compacted Ice Edge 


D. Motion 

. Compacting 


' Category A compulsory. Coding units: seaman’s watch (2x). 

Table 6: Intercoder Agreement Coefficients. Category A - Concentration. 

Coding Session Phase II Regrouped Phase II Phase III 
1 .422 .653 .666 
2 .463 .653 .603 
3 .468 .689 ; 
ee eee v 
Average Differences +.214 -.019 

' Phase III was only repeated twice. 


Concluding Remarks 

Although this case study was not directed specifically to the climatic anomaly of 1816, its 
relevance pertains to the entire time frame in which this event occurred. There is a potentially 
large volume of climatic information relating to key volcanic episodes that is in a descriptive 
format. This study provides a solution to the problem of interpreting this type of information 
reliably, an approach that in the past has been superficially addressed. Reconstructions based on 
categories that are developed from thoroughly-tested evolutionary process provide information 
that describes the reliability with which the original documents were interpreted. Furthermore, 
the results of these reliability tests must accompany the reconstruction so that the question of an 
acceptable level of reliability is relegated to the user of the reconstruction to a certain degree. 
This does not mean that any agreement coefficient should be accepted by the researcher. In the 
search for an acceptable level of reliability, an attempt should be made to balance the amount of 
information obtained against the degree of objectivity by which it was derived. In this study, a 
considerable amount of information was discarded throughout the testing procedure from the first 
phase to the final code so that sea-ice concentration was the only category to be used in the final 


River Ice and Sea Ice in the Hudson Bay Region during the Second 
Decade of the Nineteenth Century 

A.J.W. Catchpole’ 


Analysis of documentary sources in the Hudson’s Bay Company Archives has provided records 
of river- and sea-ice conditions in the Hudson Bay region during the eighteenth and nineteenth 
centuries. These include six records of dates of first-breaking and first-freezing of routes to the 
bayside trading posts. Several different validity tests have been applied to these data, and the 
results of these tests generally indicate the data are valid measures. The values of river- and sea- 
ice data in each year from 1810 through 1820 are compared with their values during the whole 
period of record. This enables the identification of years with anomalously early or late dates of 
breaking and freezing, and years with severe summer Sea ice. Evidently, exceptionally cold 
summer weather occurred in the second decade of the nineteenth century. This was not initiated 
after the eruption of Tambora in April 1815, but was first apparent in 1811 and 1812. However, 
the most severe summer cold in the decade occurred in the two years following the eruption. 


In subarctic regions the dispersal of ice in spring and freezing of water bodies in fall are 
intimately linked to weather and climatic conditions. Anomalous weather in a particular year may 
cause exceptionally early or late breaking and freezing of rivers, lakes and seas. Ice observations 
therefore figure prominently among the routine climatologic and oceanographic observations made 
by the nations fringing the poles. Another property of ice is that it occurs in forms vividly 
apparent to casual observers, and under circumstances where it can severely restrict their physical 
activities. For these reasons, informal descriptions of ice also occur prominently in written 
historical sources that contribute to the reconstruction of climates in the recent past. So it was that 
the daily journals kept by servants of the Hudson’s Bay Company graphically described the long 
anticipated breaking of the rivers in spring and their equally vital refreezing in fall. Likewise, the 
log-books of the Company’s supply ships that annually sailed the ice-congested waters of Hudson 
Strait and Hudson Bay gave frequent, detailed descriptions of the ice which imperilled their 

The Ice Records 

These sources have yielded six records of dates of first-breaking and first-freezing of the estuaries 
of rivers draining into Hudson Bay and three records of summer sea-ice severity encountered 
along portions of the sailing routes to the bayside trading posts. The various records commenced 
in the early or mid-eighteenth century, and in most cases ended in the latter part of the nineteenth 
century. Table 1 lists for each record its location, the year when the record commenced and 
ended, the total number of years in which ice data have been reconstructed and the source in 
which the reconstruction was originally published. Three stages in the seasonal development of 
river ice are dated: 

' Department of Geography, The University of Manitoba, Winnipeg, Manitoba R3T 2N2, Canada. 


1. date of first-breaking - the first day on which any evidence of breaking was observed, 
irrespective of whether or not the river remained broken thereafter; 

2. date of first partial freezing - the first day on which the river was observed to become 
partially frozen, irrespective of the spatial extent of the ice cover or its continuity thereafter; 

3. date of first complete freezing - the first day on which the entire surface of the river was 
frozen, irrespective of whether or not it remained completely frozen thereafter. 

Each of these historical records was derived from daily journals written at trading posts located 
in the estuaries of rivers draining into Hudson Bay (Figure 1). In several of these estuaries the — 
locations of the posts changed from time to time, but the records derived in the Severn, Albany, 
Moose and Eastmain estuaries are each based on single post locations. The Churchill journals 
were written both at the Old Fort, located inside the estuary, and at Fort Prince of Wales situated 
on an exposed promontory where the north shore of the estuary protrudes into Hudson Bay. The 
Churchill first-freezing data used in this paper were those reconstructed from the Old Fort 
journal, and the first-breaking data were those derived at Fort Prince of Wales. In the Hayes 
River estuary the location of York Factory was changed in 1791 to a site very close to the former 
on the north shore of the estuary. 

Table 1: Historical Records of River Ice and Sea Ice Derived from Hudson’s Bay Company Archives. 

Dates of First-Breaking of River Estuaries 

Location Limits of Years 
River of Record of Record of Record Sources 
Churchill Churchill Old Fort 1720-1866 110 Moodie and Catchpole (1975) 
Fort Prince of Wales 1731-1861 107 
Hayes York Factory 1 1715-1790 74 Moodie and Catchpole (1975) 
York Factory 2 1791-1851 45 Moodie and Catchpole (1975) 
Severn Fort Severn 1763-1939 104 Magne (1981) 
Albany Fort Albany 1722-1939 190 1722-1866 (Moodie & Catchpole 1975); 
1872-1939 (Magne 1981) 
Moose Moose Factory 1736-1871 133 Moodie and Catchpole (1975) 
Eastmain Eastmain House 1743-1939 109 Magne (1981) 


Table 1: (cont’d) 















Dates of First Partial Freezing of River Estuaries 

of Record 

Churchill Old Fort 
Fort Prince of Wales 

York Factory 1 
York Factory 2 

Fort Severn 

Fort Albany 

Moose Factory 

Eastmain House 


of Record of Record 





of Years 




Moodie and Catchpole (1975) 
Moodie and Catchpole (1975) 

Moodie and Catchpole (1975) 
Moodie and Catchpole (1975) 

Magne (1981) 

1721-1867 (Moodie & Catchpole 1975); 
1872-1938 (Magne 1981) 

Moodie and Catchpole (1975) 
Magne (1981) 

Dates of First Complete Freezing of River Estuaries 

of Record 

Churchill Old Fort 
Fort Prince of Wales 

York Factory 1 
York Factory 2 

Fort Severn 

Fort Albany 

Moose Factory 

Eastmain House 


of Record of Record 





of Years 






Moodie and Catchpole (1975) 
Moodie and Catchpole (1975) 

Moodie and Catchpole (1975) 
Moodie and Catchpole (1975) 

Magne (1981) 

1721-1864 (Moodie & Catchpole 1975); 
1872-1921 (Magne 1981) 

Moodie and Catchpole (1975) 
Magne (1981) 

Table 1: (cont'd) 

Summer Sea-Ice Severity Indices 

Limits of Years 
Region of Record of Record Sources 
Hudson Strait 1751-1889 137 1751-1870 (Catchpole and Faurer 1983); 
1871-1889 (Catchpole and Hanuta 1989) — 
Eastern Hudson Bay 1751-1870 108 Catchpole and Halpin (1987) 
Western Hudson Bay 1751-1869 111 Catchpole and Hanuta (1989) 
90°W 80° 70° 60°W 
65°N 7 65°N 
60° | | ——— | 60° 
York ly 
55° | - = ——*|55° 
"Be ees 
SS 0 200 400 
Albany Eastmain Re mactcciriecmnces 
| Moose 
| | 
: | | eee ON 
eae 90° W 80° 70° 60°W 

Figure 1: Location map showing sailing routes through Hudson Strait, across eastern Hudson Bay to 
Moose, and across western Hudson Bay to York and Churchill. 


The three sea-ice severity records refer not to point locations but to the three portions of the 
sailing-ship route (Figure 1). These records were reconstructed from descriptions of ice given in 
the supply ships’ log-books. These ice-severity indices are numerical in form but they function 
as ordinal not interval data. As such, the indices rank the years on the basis of summer-ice 
severity, but they are not numerical measures of the quantities of ice present in each summer. 
The frequency distributions of the ice indices are highly skewed, with very high proportions of 
small values and a few very large values. This property implies that the indices discriminate more 
accurately between the ranking of the few severe ice years than between that of the large number 
of moderate and light ice years. 

Quality of Ice Records 

The objective of this paper is to use the records listed in Table 1 to determine whether the river- 
and sea-ice conditions in the second decade of the nineteenth century were in any respects 
anomalous when compared with the ice conditions observed throughout the periods of record. In 
view of this objective it is pertinent to comment briefly on the quality of these historical data. 
Two aspects of the quality of climatic data derived from historical sources should be considered. 
These are the reliability of the method of derivation and the validity of the data derived. The 
reliability of the method determines the degree to which similar results will be obtained when the 
same method is applied to the same sources by the same person, or by different people with 
similar training. The validity of the data determines the degree to which the results are true 
measures of what they are intended to measure. There has been no fully comprehensive testing 
of the quality of these historical river- and sea-ice data. However, several studies have yielded 
information that bears upon their reliability and validity. 

The derivation of the breaking and freezing dates of the river estuaries (Moodie and Catchpole 
1975) included reliability testing as one of its major aspects. The test results showed that high 
degrees of reliability were obtained when dates based on direct dating categories were derived 
for places where the journals were kept. Much lower levels of reliability were obtained for dates 
based on less direct information. Marcia Faurer (this volume) is developing and applying an 
innovative approach to testing the reliability with which sea-ice data can be derived from sailing 
ships’ log-books. 

The validity of river-ice dates has been tested internally by examining the spatial homogeneity 
between similar dates derived at adjacent estuaries. These tests found high correlations between 
the dates of first-breaking at Fort Albany and Moose Factory and supported the conclusion that 
these are, therefore, true measures of the actual breaking dates in these river estuaries (Moodie 
and Catchpole 1976). Some studies have compared selected river-ice dates with tree-ring data 
derived from trees growing in the vicinity of the river estuaries. These studies were not designed 
as tests of the validity of the ice data but they do detect similarities between the trends revealed 
by tree-ring and ice data. In so doing they provide rudimentary indications of the validity of the 
ice data tested against external criteria. This approach is exemplified by a study of ice conditions 
in the Churchill River estuary conducted by Jacoby and Ulan (1982). This study used tree-ring 
data from near Churchill. It found a multiple correlation coefficient of 0.69 between tree growth 
and the date of complete freezing at Fort Prince of Wales during 1741-64. Jacoby and Ulan 
(1982) used this relationship to derive dates of complete freezing from tree-ring data in the period 
1680-1977. In his reconstruction of temperatures in the Hudson Bay region during the past three 
centuries, Guiot (1986; this volume) assembled a database including early instrumental 
temperature observations, tree-ring data and river-ice dates. 


Significant correlations were found between several of the records of the first-breaking and 
freezing of river estuaries and other records in this database (Guiot 1986, pp. 13, 19). Dates of 
first partial freezing and first complete freezing were generally found to be positively correlated 
with autumn temperatures measured at York and Churchill, whereas dates of first-breaking were 
generally negatively correlated with spring temperatures. Some of the tree-growth records were 
negatively correlated with the date of first complete freezing, and this finding is consistent with 
the results obtained by Jacoby and Ulan (1982). Lough and Fritts (1987; Lough this volume) used 
North American tree-ring data to assess the possible effects of volcanic eruptions on North 
American climate during 1602-1900. In this study they employed the mean dates of first-breaking 
and first complete freezing of the James Bay estuaries as "independent temperature records 
outside the area covered by the arid site tree-ring reconstructions." Using superposed epoch 
analysis, Lough and Fritts detected changes in ice dates following major volcanic eruptions that 
were consistent with the observed changes in tree growth. 

Wilson (1988; this volume) has derived summer thermal indices for the southeast coast of Hudson 
Bay in the nineteenth century, using a miscellany of historical evidence in the Hudson’s Bay 
Company Archives. A preliminary study of these indices shows that they may afford an indirect 
means of testing the validity of sea-ice data, in so far as anomalous summer cold in this region 
may be a result, or a cause, of severe summer ice on adjacent seas. This study involved a 
comparison between the incidence of severe ice years and negative anomalies in the thermal 
indices for May to June (Figure 2A) and May to October (Figure 2B). The May to June data 
were selected for this comparison because the ships’ log-books were not among the historical 
sources used to derive these indices. The May to October data were selected because Wilson 
(1988, p. 13) considered that the index is most accurate over the entire summer season. However, 
the sea-ice indices and May to October thermal indices are not entirely independent because the 
ships’ log-books did play a minor role as sources in the derivation of the mid-summer thermal 

Figure 2 comprises graphs of Wilson’s thermal indices upon which are superimposed vertical bars 
identifying severe ice years. In this context a severe ice year is defined as one of the years having 
the 10 highest ice indices in each of the three ice records derived for Hudson Strait, eastern 
Hudson Bay and western Hudson Bay. The 10 highest indices are based on the entire period of 
the sea-ice records, not the period 1800-70. A vertical bar on Figure 2 indicates that severe ice 
occurred in that year in Hudson Strait or in eastern or western Hudson Bay. It is judged to be 
appropriate to consider these three records together in this way and not separately. Severe ice in 
Hudson Strait could retard the entry of ships into the bay to such a degree that they would not 
encounter the bay ice in July and August, but rather in September. At this time even severe late 
summer ice is generally cleared from the bay. Furthermore, the eastern and western parts of the 
bay are not separate entities in the context of ice clearing, but rather the lateral limits of the 
waters in which the last remnants of ice tend to congregate under the influence of prevailing 
winds and currents (Danielson 1971). In years with zonal atmospheric circulation these remnants 
tend to be driven towards the east and accumulate in the sailing route to James Bay. A meridional 
atmospheric circulation permits late ice to remain in the west in the path of ships sailing to 
Churchill or York Factory. 

Figure 2 reveals a tendency for severe ice years to concur with periods having negative thermal 

indices. This is most apparent in the middle of the second decade of the century, in the late 1830s 
and in the early to mid-1840s. It is noteworthy that these are generally periods in which Wilson 


(1988, pp. 7, 8) noted the quality of the thermal indices as good to excellent. It is not appropriate 
to numerically evaluate the correlation between these data because the ice indices are ordinal not 

interval data. 


* fe 

1800 1810 1820 

‘en | 
| | E 

1830 1840 1850 1860 1870 


1800 1810 





1830 1840 1850 1860 1870 


year in which one severe ice index 

year in which two severe ice indices 

Figure 2: Thermal indices for the southeastern coast of Hudson Bay (from Wilson 1988), and years with 
severe summer ice in Hudson Strait and Hudson Bay, 1800-70. The thermal indices are 
estimates of departures from the 1941-70 normals of temperatures in the May to June (A) and 
May to October (B) periods. The quality of these indices in different time intervals was 
assessed by Wilson (1988, p. 7-8). A severe ice year is defined as a year having one of the 10 
highest ice indices, in the period 1751 to 1870, within each of the three ice records. 


Ice Conditions, 1810-20 

River- and sea-ice conditions in each year from 1810 through 1820 are evaluated in Tables 2A-C 
and 3. The data given in these tables compare the ice condition in each year with the range of 
values of that condition reconstructed over the whole period of record. In the case of the river- 
ice dates (Tables 2A-C) the comparison is made by the calculation of the parameter Z.' This 
enables the identification of years with anomalously early or late dates of breaking and freezing. 
Table 4 lists these years and distinguishes between anomalies having less than 1, 2.5 and 5% 
probabilities of occurring by chance. Table 3 gives the rank order of occurrence of each sea-ice 
severity index among the indices reconstructed for the whole period of record. Table 4 identifies 
the years in which the sea-ice index was ranked among the upper 10 values in each record. 

Fourteen of the river-ice records are designated anomalous in Table 4, and all of these are 
indicative of summer cold with significantly late-breaking and early-freezing. During this decade 
there was no occasion of early-breaking or late-freezing that produced a Z value so large that 
there was only a5% probability of its occurring by chance. The greatest anomalies, in frequency 
and amount, were those of retarded first-breaking in 1817 and 1812. In 1817 the date of first- 
breaking was anomalously late in all of the river estuaries from which a date could be obtained 
in that year. The record was interrupted in 1817 at York and Severn (Table 2A). In 1812 this 
date was anomalously late in four estuaries but not at Moose or Eastmain. Furthermore, at York 
and Severn, the 1812 anomalies exhibited 5% probabilities of occurrence by chance, whereas all 
of the 1817 anomalies exceeded this level of significance. First partial freezing was significantly 
early at Eastmain in 1817 and at Churchill in 1811. First complete freezing was early at York 
and Eastmain in 1811 and at York in 1817. 

This decade was marked by severe late-summer ice in eastern Hudson Bay in 1813 and by a 
cluster of high sea-ice indices in 1815 to 1817. This cluster included the highest ice index derived 
in Hudson Strait (1816), as well as severe ice in western Hudson Bay in 1815 and in eastern 
Hudson Bay in 1816 and 1817. 1816 provides a case in which the passage of the ships through 
Hudson Strait was so greatly delayed that they apparently entered the bay so late as to reduce 
their ability to monitor a mass of ice that persisted late in the summer within the bay. This ice 
was located in the east across the sailing route to James Bay. The ship in question (the Emerald) 
rounded Mansell Island and entered Hudson Bay on 7 September. This was 25 days later 
(standard deviation 9.7) than the mean date on which ships bound for Moose Factory entered the 
bay in the period 1751-1870. During this delayed passage to Moose in 1816, the Emerald 
encountered ice which yielded the seventh largest index (Table 3). Probably the 1816 ice in 
eastern Hudson Bay would have ranked even higher if the Emerald had sailed these waters closer 
to the normal sailing date. In 1816, the Prince of Wales sailed to York Factory. This ship also 
entered the bay on 7 September. However, it encountered no ice on its passage to the west coast 
and there is, therefore, no evidence that this ship would have encountered exceptionally late ice 
if it had sailed into the bay earlier than this late date. 

1 Z=x-p 
x=date of breaking (first partial freezing, first complete freezing) in a particular year; 
=mean date for whole period of record; 
o=standard deviation from this mean for whole period of record. 


Table 2A: Dates of First-Breaking of River Estuaries, Standard Units Z.' 


1810 +0.31 
1811 +0.72 
1812 +2.67 
1813 +0.45 
1814 0 
1815 +1.84 
1816 0 
1817 +2.40 
1818 -0.95 
1819 +0.58 
1820 -0.81 
1 Z=xy. 


Estuary (n=number of years of record) 

n= 104 



? Estuary of Churchill River at Fort Prince of Wales. 
3 Estuary of Hayes River at York Factory 2. 





Table 2B: Dates of First Partial Freezing of River Estuaries, Standard Units Z.' 


1 Z=x-p. 

Estuary (n=number of years of record) 

Churchill? York? 
n=69 n=44 
-2.20 -1.29 
-0.18 -1.15 
-0.94 - 

+0.96 = le7/s} 

- +1.29 
-0.43 - 
-0.43 +0.71 
-0.30 +1.15 

- -0.14 

? Estuary of Churchill River at the Old Fort. 
> Estuary of Hayes River at York Factory 2. 


n= 180 





n— 152 



Table 2C: Dates of First Complete Freezing of River Estuaries, Standard Units Z.! 

Estuary (n=number of years of record) 

Churchill? York? Albany Moose Eastmain 

n=95 n=39 n=178 n=117 n=97 
1810 - - - - - 
1811 - -2.20 -1.94 -1.82 -2.15 
1812 -0.94 -2.13 -0.83 -1.15 - 
1813 -0.68 - - - - 
1814 +0.60 +0.46 +1.02 +0.84 - 
1815 - -0.47 -0.83 -1.49 -0.52 
1816 - - - -0.24 +0.56 
1817 - - -1.20 0 -1.70 
1818 +0.74 +1.68 +1.30 +1.25 - 
1819 - - -1.11 - -1.34 
1820 - - -0.83 -0.16 -0.34 


? Estuary of Churchill River at the Old Fort. 
> Estuary of Hayes River at York Factory 2. 

Table 3: Summer Sea-Ice Severity Indices, Annual Ranking. 

Location (n=number of years of record) 

Eastern Western 
Hudson Strait Hudson Bay Hudson Bay 

n= 137 n=108 n=111 

1810 88 39 32 
1811 36 66 - 
1812 36 22) 32 
1813 57 2 101 
1814 12 78 101 
1815 44 44 6 
1816 1 7 101 
1817 Di 8 - 
1818 40 95 41 
1819 48 95 101 
1820 135 81 : 15 


Table 4: Incidence of Anomalous River-Ice Dates and Severe Sea Ice During 1810-20. The Probabilities 
of River-Ice Anomalies are Based on the Standard Units Z (Tables 2A-C). The sea-ice anomalies 
are the years having one of the 10 highest ice-severity indices in each of the three records. The 
fractions given compare the rank with the number of years in the record. 

o ; +t 
& 1811 1812 1813 = 1815 1816 1817 





Churchill EASTMAIN 


York York 




(East) 7/108 8/108 


Hudson Bay SEVERE 
(West) 6/111 


CHURCHILL : Less than 1% ALBANY : Less than 2.5% York : Less than 5% 


The river- and sea-ice data presented here indicate that in the second decade of the nineteenth 
century cold summer weather was not initiated after the eruption of Tambora in April 1815, but 
was first apparent in 1811 and 1812. However, this evidence does show that the most severe 
summer cold in that decade occurred in the two years following the eruption. 

The first of these cold episodes commenced in 1811 with early first partial freezing at Churchill 
and early first complete freezing at York and Eastmain. This was followed in the spring of 1812 
with late first-breaking at Churchill and Albany and with late breaking, though less delayed, at 
York and Severn. In the fall of 1812 first complete freezing occurred early at York. 

An isolated case of severe sea-ice occurred in eastern Hudson Bay in 1813, and this was followed 
by a cluster of years with severe ice in 1815 (western Hudson Bay), 1816 (Hudson Strait and 
eastern Hudson Bay) and 1817 (eastern Hudson Bay). This period culminated in late first- 
breaking in 1817 at Eastmain, Churchill, Albany and Moose. In the fall of 1817 early first partial 
freezing occurred at Eastmain. There were gaps in the historical record during both of these cold- 
summer periods, and these were most prominent in 1816 and 1817. In particular, data on first 
partial freezing and first complete freezing are unavailable for Churchill and York in 1816, and 
no river-ice data are available for York in 1817. 



Catchpole, A.J.W. and M.A. Faurer. 1983. Summer sea-ice severity in Hudson Strait, 
1751-1870. Climatic Change 5:115-139. 

Catchpole, A.J.W. and J. Halpin. 1987. Measuring summer sea-ice severity in eastern Hudson 
Bay 1751-1870. Canadian Geographer 31:233-244. 

Catchpole, A.J.W. and I. Hanuta. 1989. Severe summer ice in Hudson Strait and Hudson Bay 
following major volcanic eruptions, 1751 to 1889 A.D. Climatic Change 14:61-79. 

Danielson, E.W. 1971. Hudson Bay ice conditions. Arctic 24:90-107. 

Guiot, J. 1986. Reconstruction of temperature and pressure for the Hudson Bay Region from 
1700 to the present. Canadian Climate Centre Report No. 86-11:1-106. 

Jacoby, G.C. and L.D. Ulan. 1982. Reconstruction of past ice conditions in a Hudson Bay 
estuary using tree rings. Nature 298:637-639. 

Lough, J.M. and H.C. Fritts. 1987. An assessment of the possible effects of volcanic eruptions 
on North American climate using tree-ring data, 1602 to 1900 A.D. Climatic Change 

Magne, M.A. 1981. Two centuries of river ice dates in Hudson Bay region from historical 
sources. MA. thesis, University of Manitoba, Winnipeg. 78 pp. 

Moodie, D.W. and A.J.W. Catchpole. 1975. Environmental data from historical documents by 
content analysis: freeze-up and break-up of estuaries on Hudson Bay 1714-1871. Manitoba 
Geographical Studies 5:1-119. 

. 1976. Valid climatological data from historical sources by content analysis. Science 

Wilson, C.V. 1988. The summer season along the east coast of Hudson Bay during the nineteenth 

century. Part III. Summer thermal and wetness indices. B. The indices, 1800 to 1900. 
Canadian Climate Centre Report No. 88-3:1-42. 


The Climate of the Labrador Sea in the Spring and Summer of 1816, 
and Comparisons with Modern Analogues 

John P. Newell’ 


The wide range of natural variability in climatic conditions at the local and regional scales makes 
it necessary to examine data from as large an area as possible in order to determine the 
significance of past departures from present-day conditions. Many authors have demonstrated that 
the spring and summer of 1816 were among the coldest ever recorded in the region extending 
from the northeastern United States to Hudson Bay. Recent research on tree rings indicates that 
climate may not have been as severe in the western United States and Canada. This study 
examines proxy-climatic data for northeastern North America, extending from southeastern 
Newfoundland to Hudson Strait and including the waters of the Labrador Sea, in an effort to 
develop a more continental view of climate during this critical period. 

The sources investigated include: weather narratives from both Newfoundland and Labrador; a 
daily weather diary from eastern Newfoundland; and sea-ice records for the waters adjacent to 
Newfoundland and Labrador. The study demonstrates that, during the spring and summer of 
1816, climatic and sea-ice conditions in northern Labrador were among the most severe ever 
recorded; however, farther south in Newfoundland, conditions were by no means as severe, and 
may have been near nineteenth century normals. 

The 1816 patterns of climatic and sea-ice conditions in Newfoundland and Labrador are compared 
with recent (post-1950) patterns of temperature, precipitation and sea-ice conditions in eastern 
North America to determine if modern analogues exist. This comparison indicates that conditions 
in 1816 have no clear analogues in the recent climatic record. However, there are patterns that, 
while not as severe, do provide some indications of the nature of the circulation in 1816. These 
patterns indicate that the circulation during the summer of 1816 was similar to the present 
normals for March and April. This agrees with the July circulation pattern for 1816 presented 
by Lamb and Johnson (1966). 


The unusual character of the summer of 1816 in the Labrador Sea is demonstrated by the 
following report from the records of the Moravian Church which operated several missions along 
the Labrador coast: "The Jemima [the moravian mission ship] arrived in the river [Thames] from 
Labrador, after one of the most dangerous and fatiguing passages ever known. As in almost every 
part of Europe, so in Labrador, the elements seem to have undergone some revolution during the 
course of last summer" (Periodical Accounts, Vol. VI, p. 263). 

Modern research has demonstrated that the summer of 1816 was unusually cold in Europe 
(Manley 1974; Kelly et al. 1984; Briffa et al. 1988), eastern United States (Stommel and 
Stommel 1979; Ludlum 1966); and Hudson Bay (Wilson 1983; Catchpole 1985). Other authors 

' 34 Cornwall Crescent, St. John’s, Newfoundland A1E 1Z5, Canada. 


have demonstrated that sea-ice conditions in both Hudson Strait (Catchpole and Faurer 1985) and 
the Labrador coast (Newell 1983) were extremely severe during the summer of 1816. By 
comparison, sea-ice conditions in the East Greenland Sea (Scoresby 1820) and near Iceland 
(Lamb 1977; Ogilvie, this volume), while more severe than normal, did not reach the record 
conditions experienced in eastern North America. 

The only previous study (Lamb and Johnson 1966; Lamb, this volume) that directly considers 
climatic conditions in the Labrador Sea during 1816 is an analysis of January and July global sea- 
level pressure patterns for the years 1750 to 1962. It includes a map of July 1816 circulation over 
the North Atlantic indicating that a 1002 mb low-pressure centre was situated over the Labrador 
Sea, giving a northerly flow along the Labrador coast. This circulation pattern is more ~ 
representative of conditions in April than of the normal circulation in July. It should be noted that 
Lamb and Johnson provide maps showing that the Labrador Sea was outside the limits of reliable 
isobars until the 1870s. While the exact data used to construct their map for 1816 are not given 
in the report, other sources (Lamb and Johnson 1959, 1961) indicate that it was likely based on 
wind data from New England and possibly Greenland. 

This paper presents the results of an analysis of proxy-climatic records from areas surrounding 
the Labrador Sea (Newfoundland, Labrador, Hudson Strait and southwestern Greenland) and an 
attempt to reconstruct the atmospheric circulation pattern in this region for June 1816. In 
addition, temperature patterns over the area in June 1816 are used to select modern analogues for 
the 1816 circulation pattern. These modern analogues are then compared with the reconstructed 
circulation pattern. The study area and locations noted in the text are shown in Figure 1. 

Analysis of Historical Data 

The following brief review of the history of the study area in 1816 provides an indication of types 
of proxy-climatic data sources available. At the start of 1816, Newfoundland was in the midst 
of a financial crisis caused by the fall in fish prices after the end of the War of 1812, and in 
February 1816 a major fire struck St. John’s, the capital of the island. At this time the main 
economic activity in Newfoundland was the inshore cod fishery. The only other significant 
economic activity was the seal "fishery" carried out off the northeastern coast each spring. 
Farther north along the Labrador coast, the Moravian Church operated missions it had established 
during the late eighteenth century. These missions were supplied each spring by a mission ship 
that sailed directly to Labrador from England. At the same time the Moravians also operated a 
number of missions in southwestern Greenland that were supplied by Danish ships sailing from 
Denmark to the Greenland settlements. Whaling ships from Britain also operated off the west 
coast of Greenland each spring. The only other significant shipping activity in the study area at 
this time involved Hudson’s Bay Company ships that sailed from England to Hudson Bay each 
spring and returned in the fall. 

A review of material available in the Newfoundland Archives revealed that government 
correspondence from this period is rather limited. This is partly due to the fact that prior to 1818 
the governor was only resident in Newfoundland during the summer. The only pertinent remark 
was: "The weather during the greater part of the season [summer 1816] has been particularly 
unfavourable for the curing [the cod was dried in the sun] of fish" (Report of Fishery, December 
1816, Government Letter Book, Newfoundland Archives). This situation could result from either 
damp weather or calm weather with clear skies. Analysis of historical catch statistics for cod in 
Newfoundland waters (Forsey and Lear 1987) indicate that 1816 was a relatively good year. 




C. Chidley 

C. Farvel 




Battle Hr. 
© Belle Isle 


Figure 1: Study area. 


While no statistics on the seal catch are available for 1816, available data do not point to a bad 
year. The records indicate however a very low catch in 1817, which was attributed to severe ice 

A weather diary kept at Trinity, Newfoundland by the firm of Slade and Kelson provides the best 
information on the climate of Newfoundland in 1816. A review of the weather remarks and 
rain/snow frequencies given in the diary do not provide any evidence for cold conditions during 
the spring or summer of 1816. Analysis of the daily reports of wind for June 1816 indicate a high 
frequency from the southwest (55%) compared to present day normals for Bonavista, 
Newfoundland (less than 30%) and compared to Trinity in 1817 (48%) and 1818 (36%). A 
comparison of air temperature versus wind direction for St. John’s, based on modern data, 
indicates that the two parameters are closely linked (Figure 2). Southwest winds are clearly warm 
winds, so it was likely that southeastern Newfoundland experienced normal to above normal 
temperatures in June 1816. The typical synoptic situation giving southwest winds over 
Newfoundland in June is a ridge of high pressure pushing northward from the Bermuda High. 

In northern Labrador and Hudson Strait, Moravian records (Newell 1983) and Hudson’s Bay 
Company records (Catchpole and Faurer 1985; Teillet 1988) indicate severe ice conditions with 
considerably delayed clearing dates. Newell (1983) states that in 1816 it was " likely that the sea 
ice had not completely cleared the coast by the start of the next [ice] season". Analysis of sea-ice 
clearing in this region based on satellite imagery for 1964-74 (Crane 1978) demonstrates that late 
clearing dates are associated with an increased frequency of northerly winds. Catchpole and 
Faurer (1985), investigating sea-ice conditions in Hudson Strait during 1816 using logs from 
Hudson’s Bay Company ships, also found evidence for an increased frequency of northerly winds 
during the summer. 

Besides providing valuable information regarding the offshore ice conditions, the Moravian 
mission reports also provide some indication of the weather experienced at the stations. The 
following remarks regarding the summer of 1816 at Okkak follow a description of the severity 
of the winter: "In spring, the frost continued so severe, that we could not work in our gardens 
at the proper time, and consequently expect but a poor crop of vegetables this year, for the whole 
summer season has been cold and dry" (Periodical Accounts, Vol. VI, p. 265). The following 
reports from the Moravian missions in southwestern Greenland suggest different conditions on 
the other side of the Labrador Sea: "It rains almost incessantly, and if it even ceases for a day, 
yet the heavens are overcast...I must say that for these four months past, we have not had one 
day on which the sun has shone throughout the whole day" (Periodical Accounts, Vol. VI, 
p. 452). An analysis of conventional meteorological data collected at the Labrador Moravian 
stations in the 1880s and 1890s suggests that in June cold dry conditions are associated with north 
or northwest winds and lower air pressures; both of which would occur with a mean low-pressure 
centre to the east and lows tracking well south of the area. The wet conditions in southwestern 
Greenland indicate that this area was near or just east of the main low-pressure centre. 

In summary, the data presented indicate that during the spring and summer of 1816 a mean centre 
of low pressure was situated in the Labrador Sea with a trough extending north into Davis Strait 
(Figure 3). At the same time the main track of low-pressure systems was across southern 
Labrador and into the Labrador Sea. This pattern would give the north to northwest winds and 
cold/dry conditions in Labrador and the wet conditions in southwestern Greenland. South of the 
storm track, southeastern Newfoundland was under the influence of the Bermuda High. The 
temperature pattern for June 1816 has very cold conditions in northern Labrador and normal to 
above normal temperatures in Newfoundland. 




Air Temp C° 

30 60 90 120 150 180 210 240 270 300 330 360 

Wind Direction Deg. 
Figure 2: Air temperature versus wind direction for St. John’s, Newfoundland. Based on data for June 

1971-87, supplied by the Atmospheric Environment Service, Scientific Service Unit, 
St. John’s. 


Modern Analogues 

To provide a check on the proposed circulation pattern for 1816 and to give more detail on the 
nature of the circulation, modern analogues for the temperature pattern observed in June 1816 
were Selected, and their circulation patterns compared to that proposed. The criteria used were 
below-normal temperatures in northern Labrador and normal or above-normal temperatures in 
southeastern Newfoundland in June. Monthly temperature patterns were obtained from maps in 
Environment Canada publications (Climatic Perspectives and Monthly Record). During the 30- 
year period 1958-87, five years had June temperatures that met the criteria (1969, 1971, 1972, 
1978 and 1986; Figure 4). 

All of the years selected as analogues had below-normal June temperatures at Churchill, 
Manitoba, on the west coast of Hudson Bay. This pattern agrees well with conditions in 1816 
when temperatures at Churchill were considerably below normal (Catchpole 1985). All of these 
years except 1986 had cool to very cold conditions in central England; in fact, June 1971 and 
1972 were colder than June 1816 (Manley 1974). The opposition of temperatures in 
Newfoundland and England agrees with the Burroughs’ (1979) finding of an inverse relationship 
between temperatures in the two areas. The agreement between conditions in the five years 
mentioned above and 1816 is not as strong when conditions in New England are considered. Only 
two of the five years (1972 and 1986) had below-normal June temperatures at Boston: however, 
in all five years below-normal temperatures reached some part of New England. 

In all but one of the five years considered, the mean centre of low pressure in the North Atlantic 
was near its normal position, over or near the Labrador Sea. The exceptional year was 1972, 
when the low was south of Iceland. However, in all cases mentioned, the circulation was more 
intense than normal. Of the five years considered, the circulation patterns in 1969 and 1971 seem 
most unlike that for 1816. In these cases the surface winds in northern Labrador had a strong 
southerly component - totally unlike 1816. Perhaps the surface temperatures indicated for 
northern Labrador during these years (based on data from surrounding stations) are in error, and 
the true temperatures were higher. In the case of 1972, while the temperature pattern matches that 
for June 1816 in England, New England, Hudson Bay, Newfoundland and Labrador, the nature 
of the circulation does not fit the proposed pattern for the eastern Atlantic/western Europe sector 
as proposed by Kelly et al. (1984). In fact the circulation for June 1972 is totally different from 
the conditions that usually produce below-normal June temperatures over England (Perry 1972). 

The two remaining years (1978 and 1986) both have deeper-than-normal low-pressure centres 
over the Labrador Sea; however, in 1978 a high-pressure centre occurring over Hudson Bay was 
absent in 1986. Wilson (1985) proposed that such a high was centred over Hudson Bay during 
the summer of 1816. The 700-mb circulations in both of these years have some similarities and 
some major differences. Both years have above normal 700-mb heights southeast of 
Newfoundland. This feature also occurs during the three other years selected, and is likely related 
to above-normal temperatures in Newfoundland. Farther north the 700-mb patterns for the two 
years are different. In June 1986 there is a trough along the Labrador coast with the largest 
negative height departures over the mid-Labrador coast, whereas in June 1978 the trough is 
farther west, the greatest negative height departures being over Foxe Basin. 


B [Present Day] 

Figure 3: Mean June sea-level pressure: (A) estimated for 1816 and (B) present-day normals. 


Below normal temp. 

M.S.L. Pressure [mb.] 

Figure 4: Mean sea-level pressure (mb) and regions with below-normal air temperature for June: 1969, 
1971, 1972, 1978 and 1986. 


Analysis of Environment Canada (Atmospheric Environment Service) ice charts indicate that 
clearing dates for the Labrador coast were later than normal in both 1978 and 1986; however, 
neither year represented record conditions. Ice conditions at the end of June 1978 were more 
severe than at the end of June 1986, but the rate of retreat during the month of June was greater 
in 1978 than in 1986. Since this analysis only considered conditions in June, it is not surprising 
that ice conditions were not as severe as in 1816. Conditions earlier in the spring, and the strong 
northerly flow in July 1816 indicated by Lamb and Johnson (1966), likely played an important 
role in the exceptional 1816 ice conditions. 


Comparison of circulation and temperature patterns for June 1978 and 1986 with the proposed 
pattern for June 1816 (Figure 3) indicates that they are in general agreement. For example, all 
three maps have a deep low-pressure centre in the Labrador Sea. However, apparently the 
northerly circulation in 1816 must have been more vigorous than in 1978 or 1986 to give the 
lower temperatures reported. This would require that the low-pressure centre in the Labrador Sea 
be deeper than in either of those years. The actual pattern for June 1816 likely combined features 
of both June 1978 and 1986. This pattern also agrees with the North Atlantic circulation for July 
1816 proposed by Lamb and Johnson (1966). 

The occurrence of a circulation pattern such as the one proposed for June 1816 without outside 
forcing (such as volcanic cooling) does not seem unrealistic in light of the variability 
demonstrated in the five analogues considered in this study. Perhaps such an occurrence is 
especially likely considering that in 1816 the northern hemisphere was experiencing the last stages 
of the Little Ice Age, a period when such circulation patterns would have been more common. 
The main difficulty with this argument is that data from other sources demonstrate that conditions 
during July and August 1816 were equally unusual. A long-term data set of sea-ice conditions 
for the Labrador Sea that I am currently developing may assist in determining how the summer 
of 1816 compares with modern conditions and with other summers in the nineteenth century. 

Briffa, K.R., P.D. Jones and F.H. Schweingruber. 1988. Summer temperature patterns over 
Europe: a reconstruction from 1750 A.D. based on maximum latewood density indices of 

conifers. Quaternary Research 30:36-S2. 

Burroughs, W.J. 1979. An analysis of winter temperatures in central England and Newfoundland. 
Weather 34:19-23. 

Catchpole, A.J.W. 1985. Evidence from Hudson Bay region of severe cold in the summer of 
1816. In: Critical Periods in the Quaternary Climatic History of Northern North America. 
Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:121-146. 

Catchpole, A.J.W. and M.-A. Faurer. 1985. Ships’ logbooks, sea ice and the cold summer of 
1816 in Hudson Bay and its approaches. Arctic 38:121-128. 

Crane, R.G. 1978. Seasonal variations of sea ice extent in the Davis Strait-Labrador Sea area and 
relationships with synoptic-scale atmospheric circulation. Arctic 31:434-447. 


Forsey, R. and W.H. Lear. 1987. Historical catches and catch rates of Atlantic Cod at 
Newfoundland during 1677-1833. Department of Fisheries and Oceans, Canadian Data 
Report of Fisheries and Aquatic Sciences No. 662:1-52. 

Kelly, P.M., T.M.L. Wigley and P.D. Jones. 1984. European pressure maps for 1815-16, the 
time of the eruption of Tambora. Climate Monitor 13:76-91. 

Lamb, H.H. 1977. Climate: Present, Past and Future. Vol. 2; Climatic History and the Future. 
Methuen, London. 835 pp. 

Lamb, H.H. and A.I. Johnson. 1959. Climatic variation and observed changes in the general - 
circulation, Parts I and II. Geografiska Annaler 41:94-134. 

. 1961. Climatic variation and observed changes in the general circulation, Part III. 
Geografiska Annaler 43:363-400. 

. 1966. Secular variations of the atmospheric circulation since 1750. Great Britain, 
Meteorological Office, Geophysical Memoirs 110:1-57. 

Ludlum, D.M. 1966. Early American Winters 1604-1820. American Meteorological Society, 

Manley, G. 1974. Central England temperatures: 1659-1973. Quarterly Journal of the Royal 
Meteorological Society 100:389-405. 

Newell, J.P. 1983. Preliminary analysis of sea-ice conditions in the Labrador Sea during the 
nineteenth century. Jn: Climatic Change in Canada 3. C.R. Harington (ed.). Syllogeus 

Perry, A.H. 1972. June 1972 - the coldest June of the century. Weather 27:418-422. 

Scoresby, W., Jr. 1820. An Account of the Arctic Regions with a History and Description of the 
Northern Whale-Fishery. Reprinted in 1969 by Augustus M. Kelley, Publishers, New 

Stommel, H. and E. Stommel. 1979. The year without a summer. Scientific American 

Teillet, J.V. 1988. A reconstruction of summer sea ice conditions in the Labrador Sea using 
Hudson’s Bay Company ships’ log-books, 1751-1870. Unpublished M.A. thesis, 
University of Manitoba, Winnipeg. 161 pp. 

Wilson, C.V. 1983. The summer season along the east coast of Hudson Bay during the nineteenth 
century. Part II: The Little Ice Age on eastern Hudson Bay: summers at Great Whale, Fort 
George, Eastmain, 1814-1821. Canadian Climate Centre, Downsview, Report No. 83-9. 

. 1985. The Little Ice Age on eastern Hudson/James Bay: the summer weather and climate 
at Great Whale, Fort George and Eastmain, 1814 to 1821, as derived from Hudson’s Bay 
Company records. In: Critical Periods in the Quaternary Climatic History of Northern 
North America. Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 


Spatial Patterns of Tree-Growth Anomalies from the North American 
Boreal Treeline in the Early 1800s, Including the Year 1816 

Gordon C. Jacoby, Jr.’ and Rosanne D’ Arrigo’ 


Tree-growth anomalies based on 24 temperature-sensitive white spruce chronologies from boreal 
treeline sites in North America are mapped and analyzed for the interval 1805-24. This interval 
includes the volcanic eruption of Tambora in 1815 and the unusual "year without a summer" in 
1816. The first few decades of the 1800s also are concurrent with a series of low-amplitude 
cycles in sunspot number which have been suggested as contributing to unusually cooler 
conditions during this time. It is inferred from the tree-ring data that climatic changes following 
the Tambora eruption influenced the North American boreal forest in different areas at different 
times from 1816 to 1818, with the coldest regional temperatures appearing to have occurred in 
the year 1816 in easternmost Canada. The series of anomaly maps provided here help to clarify 
the spatial patterns of climatic changes in remote northern regions during this extreme cooling 


The climatically unusual "year of no summer" (1816) was primarily documented as such by 
observers in Europe and eastern North America (Landsberg and Albert 1974; Stommel and 
Stommel 1983; Stothers 1984; Briffa et al. 1988). Cold air masses invaded most areas of Europe 
and the settled, eastern regions of North America (Stommel and Stommel 1983; Stothers 1984; 
Briffa et al. 1988). There is little documentation on weather variations for western North America 
about the time of the Tambora eruption (1815), and much of the documentation on weather 
variations for other parts of the world is only recently being brought into full consideration (e.g., 
Legrand and Delmas 1987, concerning evidence of the Tambora eruption in Antarctic ice-core 
data; also other papers in this volume). Overall, little is known regarding the regional-scale 
climatic variations following many volcanic eruptions (including Tambora), and it is likely that 
hemispheric-scale studies demonstrating a general cooling effect may obscure warming in some 
areas (Lough and Fritts 1987) that may result from changes in large-scale atmospheric dynamics 
following major eruptions (Hansen et al. 1978; Schneider 1983). 

It has been hypothesized that volcanism can strongly influence climate, causing cooler 
temperatures (e.g., Lamb 1970; Mass and Schneider 1977; Sear et al. 1987; Bradley 1988). The 
mechanism is not thoroughly understood but the common theory is that stratospheric sulphate 
particles partially reflect and absorb incoming radiation, heating the stratosphere. This heat does 
not reach the troposphere, which then becomes cooler (e.g., Hansen ef al. 1978). Although 
empirical modeling studies (Hansen et al. 1978) and superposed epoch analyses (Mass and 
Schneider 1977; Sear et al. 1987; Skinner, this volume) indicate a cooling of a few tenths of an 

'  Tree-Ring Laboratory, Lamont-Doherty Geological Observatory of Columbia University, Palisades, New York 

10964, U.S.A. 


degree following major eruptions, this cooling is within the level of natural climatic variability. 
Hence a link cannot be unequivocally proven, but there is strong evidence for a cause and effect 
relationship (Sear et al. 1987). 

The eruption of Tambora in April 1815 was one of the greatest volcanic events of recent centuries 
(Simkin et al. 1981) . Because studies indicate (Sear et al. 1987) that southern hemisphere 
eruptions would cool northern hemisphere temperatures after a lag of about six months to a year, 
a response to this event at northern mid- to high-latitudes would not be expected to occur until 
late 1815-16. An unusually cold summer in 1816 is attributed by many to the effects of this event 
but the overall climatic anomalies of the period are more complex than a single event-response 
phenomenon. One phenomenon to be considered in the complexity is that the early 1800s were 
notable for a reduction in solar sunspot amplitudes, possibly reflecting solar irradiance changes 
(Lean and Foukal 1988; Kuhn et al. 1988) that may have contributed to a cooling during the 
early decades of the 1800s (Eddy 1977, this volume). 

Perspective of This Study 

The method used here to examine the climate of 1816 is the study of old-aged trees growing 
along the northern boreal forest zone of North America. All of the data are from white spruce 
[Picea glauca (Moench) Voss] near the forest-tundra transition zone. These trees primarily 
respond to summer temperature, with a secondary response to fall and spring temperature-related 
conditions (Jacoby and D’Arrigo, in press; Scott et al. 1988; Jacoby and Ulan 1982; Cropper 
1982; Jacoby and Cook 1981; Garfinkel and Brubaker 1980). The average position of the Polar 
Front in summer largely coincides with the location of the northern treeline (Bryson 1966). Tree 
growth in this region can therefore be expected to record frontal shifts, extensive outbreaks of 
polar air masses and circulation changes influencing thermal conditions (e.g., Scott et al. 1988). 
Thus tree-ring data can provide useful information on climatic response following the 1815 
eruption (and the early 1800s in general) in the northern boreal region, including the western 
region of North America where other data sources are scarce. 

We have examined 24 time-series of absolutely dated tree-ring width indices (chronologies) 
throughout the region. Each time-series is usually based on about 10 trees with multiple cores 
(radii) from each tree. Some of these chronologies contain low-frequency response to climate, 
whereas others only preserve higher-frequency response. For compatibility and intercomparison, 
all chronologies were prewhitened (to remove low-frequency variation) and normalized. This 
procedure is appropriate since, in this case, we are evaluating variations in year-to-year response 
over a period of a few decades. Maps (Figures 1, 2) display the departures from the mean for 
each chronology location during the early 1800s (1805-24). 

Distribution of Anomalies 

For 1816, Figure 1 shows substantial negative departures (reflecting reduced radial growth/colder 
temperatures) in eastern Canada. These are the greatest anomalies for the period under review 
(1805-24). However, the rest of Canada and Alaska shows no such severe cooling. Northern 
Alaska was fairly cold (as in some other years), but central and western Canada are close to 
average or above for the year. The colder eastern region agrees well with reports and records 
from eastern Canada and the United States. For example in the northern region, Catchpole and 
Faurer (1983) demonstrate that the duration of westward passage of Hudson’s Bay Company ships 
was the longest (54 days compared to a mean of 17.7 days) of the entire 1751-1870 record, 


representing severe sea-ice conditions. The authors (see also Catchpole, this volume) suggest that 
these conditions could be explained by enhanced meridional flow of arctic air masses over eastern 
North America at this time. Records from the eastern United States, do not show a continuously 
cold summer (Stommel and Stommel 1983; Baron, this volume). There were three distinct 
outbreaks of extremely cold air from the North at different times during the summer. These 
outbreaks had serious negative effects on food crops during the growing season (Stommel and 
Stommel 1983). Such records appear to support the theory of increased Arctic air flow over this 
region in 1816. Schneider (1983) suggests that since the cooling (about 3°C) would not have been 
sufficient to explain the documented frosts that occurred, a dip in the Jetstream and blocking of 
the mid-latitude westerlies could have contributed to the adverse conditions. He suggests that 
conditions to the west (about one-half wavelength away) of eastern North America would have 
been unusually warm if this had been the case. Our results support this contention since 
conditions in central Canada, although not unusually warm, do not demonstrate the pronounced 
cooling found in the east (Figure 1). 

Figure 1: Tree-growth anomaly map for 1816. The growth departures are based on prewhitened and 
normalized tree-ring width indices for 24 white spruce chronologies from near the boreal 
treeline of North America. 

To place this year in context, we review the years preceding and after 1816 beginning with 1805 
(Figure 2). The three years of 1805-07 show little in the way of extreme cold temperatures. 
Except for southeastern Alaska, most other regions are near or above normal, and eastern Canada 
is substantially above normal. Then in 1808 colder temperatures prevailed in the Hudson Bay 
region, and Alaska was warm. In 1809 the Hudson Bay region was less cold but Alaska became 
cold. The distribution of regions of warmer or cooler temperatures correspond roughly to the 
configuration of the longwave pattern in the atmosphere. There is approximately one wavelength 
across the North American quadrant for a four-wave pattern, western Alaska to eastern Canada 
being slightly over 90° of longitude (Chang 1972). 


“1.0. 6 -0.7 0.5 

Figure 2: Tree-growth anomaly maps for 1805-24. The growth departures are derived as in Figure 1. 


Figure 2 (cont’d): 


0o0 08 

“1.1. 2 0.0 

01. 204/05 


Figure 2 (cont’d): 


Figure 2 (cont’d): 

-0.8. »-0.6| -0.7. 

1.6» 0.8 | 0.6 


During 1810 through 1812 the main features of the maps are a cooling over eastern Alaska, the 
Yukon Territory and the western Northwest Territories, and in Labrador an alternating 
warm-cool-warm sequence. An indication of very warm conditions in 1813 over the Northwest 
Territories is followed by a reversal to quite cold conditions for Alaska, especially southeastern 
Alaska, and all of western Canada in 1814 accompanied by a warming in Labrador. The cooler 
conditions continue for the western region in 1815. Northern Alaska is fairly cold in 1816, but 
the most severe cold is restricted to eastern Canada (see also Figure 1). As noted above, central 
and western Canada are not unusually cold during 1816. More severe cold does not reach western 
Canada and eastern Alaska until 1817 when the anomalies are more negative than other years of 
the period, although 1809 is quite cold. The eastern region is still cold but recovering toward 
normal. In 1818, the coldest conditions occur in the Northwest Territories. Again these are the ~ 
greatest negative anomalies for this area, although 1814 and 1821-22 are also cold. By 1819, 
almost all extreme negative anomalies are gone from the map region. 

Extreme western Alaska and eastern Canada are warm in 1820 but cold pervades much of the 
entire map region during 1821 and 1822, except for Labrador in 1822. Alaska warms in 1823 
but there is a return to cold temperatures in 1824 in both western Canada/Alaska and in eastern 

In summary, there were significantly cold conditions in some northern areas before 1816. After 
the volcanic event, extreme cold affected all of the map region at different times until 1818. Cold 
temperatures pervaded some areas after 1818, and 1824 was fairly cool throughout most of the 

Discussion and Conclusions 

Our results show the spatial patterns of tree-growth anomalies from the North American boreal 
treeline during the anomalous period of the early 1800s, with an emphasis on the year 1816. The 
data provide added spatial coverage of western Canada and Alaska in relation to the climatic 
response following the Tambora (1815) event. In agreement with other studies, apparently the 
unusual cold in eastern North America may in part have resulted from meridional flow of cold 
Arctic air across this region. By contrast, conditions in western and central Canada (about a half 
wavelength to the west) were moderate in 1816 (Schneider 1983). This reflects a possible shift 
in atmospheric circulation which may or may not be directly linked to the volcanic event. 

Records of climatic conditions in the early 1800s in other regions of the globe are rather 
fragmentary, as discussed, and the cooling in the early 1800s was probably not synchronous 
globally (see the Workshop section, this volume). Figure 3 shows a reconstruction of northern 
hemisphere annual temperatures (from Jacoby and D’Arrigo 1988) based on North American 
boreal tree-ring data, indicating a cooling during this interval that persisted for several decades. 
On a more regional scale, the cooling in Europe is well documented (Stommel and Stommel 
1983; Briffa et al. 1988). Specifically, a sharp lowering of temperature is seen in England and 
central Europe from 1812-20 in reconstructed summer temperatures based on tree-ring density 
data (Briffa et al. 1988). Records from China and Japan (Stommel and Stommel 1983) do not 
indicate unusually cold conditions’. A detailed compilation of spatial temperature data during this 
time interval is clearly needed. 

' But see Zhang et al. and Huang, this volume (editor). 





1670 1710 1750 1790 1830 1870 1910 1950 1990 


Figure 3: Reconstruction of annual northern hemisphere temperatures from 1671 to 1973 based on high 
latitude tree-ring data from North America. Temperature departures from 1974-87 from Hansen 
and Lebedeff, 1987, 1988 [see Jacoby and D’ Arrigo (in press)]. Note the abrupt cooling in the 
early 1800s. 

The detection of a direct cause and effect signal due to volcanism is difficult, in part due to the 
influence of other forcing functions on the climatic system. These include unusual (diminished) 
solar fluctuations that also occurred in the early 1800s (Eddy 1977). El Nifio events occur on the 
same time scale as volcanic eruptions (largely high frequency) and can obscure their signal, as 
can random climatic variations (Robock 1981). Modeling studies (e.g., Gilliland and Schneider 
1984; Robock 1981) show good agreement between model estimates (based on volcanic indices) 
and actual temperature data but other forcings must also be considered. Finally there are many 
complicating factors for individual eruptions [e.g., season and latitude of eruption, height and 
chemistry of ejecta, state of atmospheric circulation at time of eruption (Lamb 1970; Lough and 
Fritts 1987)] which complicate attempts to detect a common event-response pattern. 
Improvements in understanding volcanic forcings are necessary for isolating effects of other 
forcings such as CQ). 

The oft-applied term "year of no summer" for 1816 is obviously a misnomer in the context of 
the northern boreal forests of North America. To understand climatic change in Canada and the 
rest of North America, it is necessary to move away from this oversimplification and study spatial 
and temporal differences and dynamics of the early 1800s, as these decades are a time of 


substantial climatic variation. Here we have provided a series of maps from the early 1800s that 
help clarify the spatial patterns of climate at remote high-northern latitudes during this interval, 
and which may be useful in determining causes of and responses to such extreme climatic events. 


This research was supported by the Climate Dynamics Division of the National Science 
Foundation, under grants ATM85-15290 and ATM87-16630. We thank J. Hayes and 
W. Ruddiman for helpful reviews, and the Canadian Forestry and Atmospheric Environment 
services for technical assistance. Lamont-Doherty Geological Observatory Contribution No. 4566. 


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Gilliland, R.L. and S.H. Schneider. 1984. Volcanic, CO, and solar forcing of northern and 
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Jacoby, G.C. Jr. and E.R. Cook. 1981. Past temperature variations inferred from a 400-year 
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Jacoby, G.C. Jr. and R. D’Arrigo. 1988. Reconstructed northern hemisphere annual temperature 
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Early Nineteenth-Century Tree-Ring Series from Treeline Sites 
in the Middle Canadian Rockies 

B.H. Luckman! and M.E. Colenutt' 


Preliminary data from tree-ring series at treeline sites in the Canadian Rockies are evaluated for 
evidence of anomalies associated with the 1815 eruption of Tambora. Three maximum density . 
and nine ring-width chronologies (six Picea engelmannii, one each for Abies lasiocarpa, Larix 
lyallii and Pinus albicaulis) are presented, covering 1780-1860. The absence of narrow or light 
latewood marker-rings associated with 1816 or 1817 indicates that there is no distinctive tree-ring 
signal associated with the Tambora event at these sites. Most records do however contain a sharp 
decrease in ring widths during the 1810-20 decade, similar to that reported from latitudinal 
treeline sites elsewhere in North America, and which appears to be associated with an abrupt 
deterioration of climate initiated several years prior to the Tambora eruption. 


Evaluation of the spatial extent of climatic anomalies associated with the eruption of Tambora in 
1815 requires assessment of proxy-data series throughout North America. The annual resolution 
of tree-ring series, combined with strong relationships between ring characteristics and climate, 
potentially provide a powerful tool to accomplish this goal. Here we present data from 
preliminary tree-ring chronologies for 1780-1860 at several treeline sites in and adjacent to Banff 
and Jasper national parks in the Canadian Rocky Mountains (Figure 1). These data are examined 
to see whether significant anomalies are present in the 1815-17 period that could be attributed to 
climatic effects associated with the Tambora eruption. 

The principal direct climatic effect associated with volcanic eruptions is a reduction in solar 
radiation received at the surface because of stratospheric dust veils (Lamb 1970). This often 
results in cooler summers, and the spatial extent and severity of this effect depends on the 
magnitude, timing and nature of the eruption. LaMarche and Hirschboeck (1984) have 
demonstrated a strong relationship between the presence of frost rings in the Bristlecone pine 
chronology from treeline sites in the White Mountains of California and major volcanic eruptions. 
These data include a frost-ring date of 1626 B.C. for the eruption of Santorini in Greece (which 
destroyed the Minoan civilization of Crete, and probably had global climatic effects). Baillie and 
Munroe (1988) show that Irish bog oaks had very narrow rings in the 1620s (B.C.), which appear 
to confirm this result. In Canada, Filion et al. (1986) have shown that light latewood rings from 
black spruce chronologies in northern Quebec correspond with periods 0-2 years after major 
volcanic eruptions. In these records the 1816-17 rings have light latewood in 75% of the series 
studied, and 1784 (the year following the Laki eruption) is also a prominent marker ring. Parker 
(1985) and Jacoby et al. (1988) also note the exceptional nature of the 1816 and 1817 rings in 
white-spruce chronologies on the eastern shores of Hudson Bay. The severe climate of these two 
summers is amply demonstrated by several papers in this volume. 

' Department of Geography, University of Western Ontario, London, Ontario N6A 5C2, Canada. 


Baal Sal 


, Y yy 

Lake Louisel//, 

Larch Valley */// 




Figure 1: Location of the main study sites. 

Parker (1985) evaluated selected tree-ring series from western and central Canada to determine 
whether he could detect a signal associated with the eruptions of Tambora or Krakatau (1888). 
He used ring-width and densitometric data from 135 trees (four different species) at 15 sites 
between Vancouver Island and Hudson Bay. The data were aggregated into six regional 
chronologies, and indexed data were used to compare the eruption year with groups of three years 
before and after the eruption. Only one site, Cri Lake on Hudson Bay, showed a significant 
growth reduction following the Tambora eruption. 

These results suggest that, under certain conditions, a volcanic signal can be detected via its 
influence on climate, and thereby on tree-ring characteristics. Many authors have demonstrated 
strong relationships between tree-ring width or density series and summer temperatures - 
particularly at treeline sites (e.g., Parker and Henoch 1971; Luckman et al. 1985; Jacoby and 
Cook 1981; Jacoby et al. 1988). It would be anticipated, therefore, that trees at these sites would 
be particularly sensitive to reductions in summer insolation, and therefore most likely to record 
evidence of dust-veil-related volcanic effects. As most of the montane sites used by Parker (1985) 
are well below treeline, we decided to evaluate treeline records from the Rockies to see whether 
the 1815-17 record contained any distinctive signal that could be attributed to the effects of the 
Tambora eruption. 


Sample Sites 

Nine preliminary living-tree ring-width chronologies are available from our tree-ring studies in 
the Canadian Rockies (Table 1, Figure 1): eight are from treeline sites and five (Robson, 
Bennington and Icefields/Athabasca sites) are adjacent to Little Ice Age terminal moraines. Six 
of these chronologies utilize Engelmann spruce (Picea engelmannii) because it is the most 
ubiquitous, long-lived tree at treeline in this area. Single chronologies for alpine larch (Larix 
lyallii), alpine fir (Abies lasiocarpa) and whitebark pine (Pinus albicaulis) are also used. 
Tree-ring densitometric data are also available for three of these sites. 

The Robson site (Figure 2) is an isolated stand of spruce on a low bedrock knoll overlooking an — 
inactive outwash fan from Robson Glacier. At its Little Ice Age maximum position, Robson 
Glacier advanced against the upvalley side of the knoll and built a terminal moraine along its 
crest. Heusser (1956) estimated the date of formation of the three outermost moraines of Robson 
Glacier as 1787, 1801 and 1861 based on tree-ring sampling and allowing a 12-year ecesis 
interval. The oldest tree in the stand outside the moraine is just over 400 years old, i.e., it 
predates the maximum glacier advance by about 200 years. The trees were sampled in 1981 and 
1983. Preliminary results are given by Watson (1983): the results presented here use both data 


Figure 2: The Robson Glacier site, view east 
from Adolphus Lake, Alberta 
(foreground) toward Rearguard 
Mountain and Mount Robson 
(snow-covered, top right). The 
sampled stand (a) is visible with the 
lighter-toned Little Ice Age moraine 
complex of the Robson Glacier 
extending from left to right across 
the middle ground behind the trees. 


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The Bennington site is an open grown, almost pure stand of whitebark pine growing on a coarse 
talus and bedrock slope overlooking the lateral moraine of Bennington Glacier (Figure 3) which 
is dated to about 1700 and 1825 by dendrochronology (MacCarthy 1985). This stand contains a 
number of very old trees, the oldest of which has a pith date of 1112 A.D. at breast height, and 
is thought to be the oldest whitebark pine in Canada (Luckman et al. 1984). Chronology 
development at this site is incomplete, and data given are for the four trees for which ring width 
and densitometric data are presently available for the 1780-1860 interval. The spruce stand at this 
site is on the lower slope, slightly downvalley of the area shown in Figure 3. The tree-ring series 
from this site show high tree-to-tree variability due to rockfall disturbance of the site (Watson 
1983). The results are included here solely for comparison with spruce chronologies at other 


The Athadome (Figure 4) and Icefield (Figure 5) sites are both adjacent to the Athabasca Glacier 
on opposite sides of the Sunwapta Valley, less than a kilometre apart. The Athadome site has a 
unique microclimate because it lies between the lateral moraines of Athabasca and Dome glaciers 
and was almost completely surrounded by and below the level of the adjacent ice surface about 
1714 (Heusser 1956) and between approximately 1840 and 1920 (Luckman 1988). By contrast, 
the Icefield site is a well drained lower valley side slope just beyond the outer limits of the 
Athabasca Glacier. Both chronologies are Engelmann spruce, but sampling in 1980 and 1981 
(Luckman 1982) indicated that some trees at the Icefield site were considerably older. Intensive 
sampling at this site in 1982 provided the present chronology (Jozsa et al. 1983), which is based 
on trees with a mean age of over 500 years (Table 2) and includes the oldest known Engelmann 
spruce (Luckman ef al. 1984). 

The Lake Louise site is the only non-treeline site presented here. It occurs in the lower subalpine 
forest about 300 m below treeline on a valley side bench overlooking Lake Louise townsite 
(Hamilton 1984). This site was the closest Engelmann spruce stand to the meteorological station 
at Lake Louise, and was used to explore climatic tree-ring relationships for this species (Luckman 
et al. 1985). The Larch Valley site (Figure 6) is at treeline, some 10 km south of Lake Louise. 
It is about 2 km from the Wenkchemna Glacier, and considerably above it on a broad valley side 
bench overlooking the main valley. Chronologies were developed for three species in the same 
stand at this site because of difficulties in crossdating the larch record which has several periods 
with very tight or missing rings (Colenutt 1988). This larch chronology is the best-replicated and 
most sensitive (mean sensitivity 0.38) of those discussed here. The Larch Valley and Lake Louise 
chronologies are also less likely to show local climatic effects from adjacent glaciers than the 
other chronologies reported here. 

Chronologies for most sites were developed by standard methods using the Laboratory of 
Tree-Ring Research (Tucson) programs INDEX and SUMAC (Graybill 1982). However, 
chronologies for the Icefield and Lake Louise sites were developed by Forintek, Vancouver using 
a 99-year running mean to remove the growth trend (Parker et al. 1981). Therefore some of the 
longer-frequency trends in these chronologies have been removed resulting in a lower amplitude 
of response (Luckman ef al. 1985). 

The results from the nine indexed ring-width chronologies are shown in Figure 7, and high-pass 

filter data (Fritts 1976) from these series are presented in Figure 8. The indexed values for 
1810-20 are listed in Table 2 with some summary statistics for the chronologies used. These 


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Figure 3; The Bennington Pine site, view north across the north lateral of Bennington Glacier (July 1986); 
the main sampled area in the centre of the photo contains many standing snags and trees over 

500 years old. The spruce and fir flanking the moraine are much younger in age. The ridgecrest 
is about 1000 m above the valley floor. 

Figure 4: View south across the forefield of the Athabasca Glacier to the lateral moraines of Athabasca 
Glacier (left) and Dome Glacier (right). The Athadome site is near the small group of trees at 
(a), the Icefields site is visible in the foreground, north of the Icefields Parkway. 

Paf 74 

Figure 5: View west across the forefield of the Athabasca Glacier to the Icefields site (area b). Note the 

well-marked trimline denoting the Little Ice Age limit (about 1842 A.D.) of the Athabasca 
Glacier against the slope. 

Figure 6: Part of the main stand sampled at Larch Valley. View north toward Sentinel Pass with Mount 
Temple (right). 


chronologies show considerable differences in the amplitude and nature of tree-ring response at 
these sites. This may be attributed to a number of factors such as site-to-site differences in 
microclimate or other site factors, differences in response between species (e.g., the Larch Valley 
sites), differences in vigour between sites due to age (e.g., Athadome and Icefield), differences 
in standardization procedure and length of record used to derive indices (compare for example, 
Larch Valley fir and Athadome spruce). Despite this diversity, a number of common elements 
also occur. Except for the two most northerly spruce chronologies, 1799 and 1824 are readily 
identifiable (Figures 7 and 8) as conspicuous, narrow marker rings that bracket the period of 
particular interest here [in fact, the 1824 ring was missing in 27 of the 35 larch cores measured 
at Larch Valley, (Colenutt 1988)]. Generally the records (Figure 7) can be divided into three 
parts: a period of relatively high growth, particularly between about 1790 and 1810; a period of - 
declining growth, usually from about 1810 to 1820 or 1830; and a period of low growth or 
general recovery thereafter. The relative intensity, timing and magnitude of the decline varies 
between sites but it is particularly marked in the Larch Valley, Lake Louise and Bennington pine 
chronologies. At Bennington the oldest pine is not included in this chronology because, following 
the sharp decline in ring width in the early nineteenth century, the post-1820 rings are too narrow 
to measure accurately with densitometry. 

Several authors, particularly Jacoby and coworkers (Jacoby et al. 1985, 1988; Jacoby and 
D’ Arrigo 1989; Ivanciu and Jacoby 1988) have reported an abrupt cooling in the early 1800s 
based on high-latitude North American tree-ring series and other data. This decline is shown to 
some extent by all of the alpine treeline chronologies reported here. At most sites 1816 occurs 
in the middle, or at the end of, this period, and is not a marked departure from the trend. 
Detailed examination of the index values of these chronologies (Table 2, Figure 8) shows that 
only one of the nine chronologies (Larch Valley spruce) has a significantly narrower ring in 
1816. Two others have local minimum values in 1816, but these have similar values to preceding 
rings in 1814 (Lake Louise) and 1815 (Larch Valley larch). Based on these data, although 1816 
is often represented by a narrow ring, the 1814 or 1815 rings are narrower - a fact that cannot 
be attributed to the Tambora event. It is not possible, therefore, to detect a marked decline in 
growth in 1816 from the ring-width records at these sites. 

Filion et al. (1986), Parker (1985), Jacoby et al. (1988) in northern Quebec and Jones et al. 
(1988) in Europe report that the 1816 tree ring is distinctive because of its light latewood and low 
maximum-density values. Figure 9 shows the available (three) maximum-density indexed 
chronologies for the sites previously discussed. Although 1813 appears to be a significant marker 
ring, none of these three chronologies show light marker rings associated with 1815, 1816 or 
1817. In fact, 1816 appears to have a greater maximum density than adjacent years at these sites 
suggesting that, if anything, conditions may have been a little warmer than adjacent years (Parker 
and Henoch 1971; Luckman et al. 1985). 

Several of the papers in this volume draw attention to the possible effects of the 1783 Laki 
eruption and, in preparing this paper, the diagrams were extended to 1780 to include this period. 
The data (Figures 7, 8) show considerable variability in the 1780s but Table 3 indicates that 1784 
or 1785 is the narrowest ring for the 1780-89 decade in seven of the nine chronologies (Larch 
Valley spruce and fir chronologies have slightly lower values in 1782). 1784 is narrower than 
1783 in all chronologies and, except for the two northernmost spruce chronologies, this decrease 
is marked (7-44%). However, the relative widths of tree rings representing 1784 and 1785 are 
inconsistent: at four sites 1785 is much narrower; two sites have 1784 significantly narrower 
(including Larch Valley larch, which has a missing ring); and the indexed values are similar at 



Lake Louise 

Larch Valley 

_| Larch Valley 


Icefields|'\ aly 



1800 1820 1840 1860 1780 1800 1820 1840 1860 

Figure 7: Ring-width chronologies (1780-1860) for nine sites in the Canadian Rockies. The ring-width 

series are standardized to a mean of 1.0 over the entire period of record (200-600 years; 
Table 1) and are plotted at the same scale. The lighter line is the chronology; the thicker line 
is a 13-year low-pass filter (see Fritts 1976). 


Bennington Pine 

Bennington Spruce 

Robson Spruce 

Icefield Spruce 

Athadome Spruce 

Lake Louise Spruce 

Larch Valley Spruce 

Larch Valley Fir 

Larch Valley Larch 

1780 1800 1820 1840 1860 

Figure 8: 13-year high-pass filter of ring-width chronologies for nine Rocky Mountain tree-ring sites. 
These data are standardized and plotted at the same scale. These data correspond to the 
deviations from the low-pass filter curve (Figure 7). 1799 and 1824 are significant narrow 
marker rings at most sites. 1783 is the date of eruption of Laki in Iceland. 



| Icefields 


wa | Lake Louise 

1780 1800 1820 1840 1860 
Figure 9: Standardized maximum density (MXD) chronologies for three sites in the Canadian Rockies. 

All are plotted at the same scale. The thin line is the annual indexed value; the thicker line is 
a 13-year low-pass filter (Fritts 1976). The shaded year is 1813. 


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the three other sites. Generally, these data indicate that the two years following the 1783 Laki 
eruption are considerably narrower, but the pattern is not consistent enough to use 1784 or 1785 
as a marker ring (Figure 8). 


In this paper we have examined nine ring-width chronologies and three maximum-density 
chronologies from four different species at treeline sites in the central Canadian Rocky 
Mountains. Although the 1810-20 decade showed a marked decrease in ring-width at all sites, 
probably as a result of climatic cooling, there is no indication of a marker (narrow) ring 
associated with 1816 or 1817 in the years following the Tambora eruption. Examination of the - 
three maximum-density series available for these sites indicated no light latewood rings in 1815, 
1816 or 1817. It would therefore appear that, unlike sites east of Hudson Bay, the treeline sites 
we have examined have no distinctive tree-ring signal to suggest significantly poorer growth 
conditions in 1816 or 1817. A major growth decline is identified at all sites during 1810-20 but, 
as that decade begins some years prior to the Tambora eruption, that eruption cannot be its 
principal cause. This significant period of declining ring-width has been identified elsewhere in 
North America, and reflects the most abrupt deterioration in climatic conditions during the last 
few centuries. 


We thank: the Natural Sciences and Engineering Research Council of Canada for support of this 
research; Parks Canada and Mount Robson Provincial Park staff for permission to carry out 
research at these sites; L. Jozsa, Forintek Canada Corporation, for assistance in the field and in 
processing density data; F.F. Dalley (1980), G. Frazer (1981, 1983), S. Ulansky (1982), 
J. Hamilton (1983-84), D.C. Luckman (1985-87), D. McCarthy (1986), R. and S. Colenutt 
(1987) for coring assistance; and M.I. Johnson, H. Watson, K. Harding, B. Schaus, J. Hamilton 
and G. Frazer for ring-width measurements. 


Baillie, M.G.L. and M.A.R. Munroe. 1988. Irish tree rings, Santorini and volcanic dust veils. 
Nature 332:344-346. 

Colenutt, M.E. 1988. Dendrochronological studies in Larch Valley, Alberta. B.Sc. thesis, 
Geography Department, University of Western Ontario, London, Ontario. 123 p. 

Filion, L., S. Payette, L. Gauthier and Y. Boutin. 1986. Light rings in subarctic conifers as a 
dendrochronological tool. Quaternary Research 26:272-279. 

Fritts, H.C. 1976. Tree Rings and Climate. Academic Press, New York. 
Graybill, D.A. 1982. Chronology development and analysis. In: Climate from Tree Rings, 

M.K. Hughes, P.M. Kelly, J.R. Pilcher and V.C. LaMarche Jr. (eds.). Cambridge 
University Press, Cambridge. pp. 21-28. 


Hamilton, J.P. 1984. The use of densitometric tree-ring data as proxy for climate at Lake Louise, 
Alberta. B.A. thesis, Geography, University of Western Ontario, London, Ontario. 116 

Heusser, C.J. 1956. Postglacial environments in the Canadian Rocky Mountains. Ecological 
Monographs 26:263-302. 

Ivanciu, I.S. and G.C. Jacoby. 1988. An abrupt climatic cooling in the early 1800s as evidenced 
by high-latitude tree-ring data. In: The year without a summer? Climate in 1816. An 
International Meeting Sponsored by the National Museum of Natural Sciences, Ottawa, 
1986. Abstracts. p. 29. 

Jacoby, G.C. and E.R. Cook. 1981. Past temperature information inferred from a 400-year 
tree-ring chronology from Yukon Territory, Canada. Arctic and Alpine Research 

Jacoby, G.C. and R. D’Arrigo. 1989. Reconstructed northern hemisphere annual temperature 
since 1670 based on high-latitude tree-ring data from North America. Climatic Change 

Jacoby, G.C., E. Cook and L.D. Ulan. 1985. Reconstructed summer degree days in central 
Alaska and northwestern Canada since 1524. Quaternary Research 23:18-26 

Jacoby, G.C., I.S. Ivanciu and L.D. Ulan. 1988. A 263-year record of summer temperature for 
northern Quebec reconstructed from tree-ring data and evidence of a major climatic shift 
in the early 1800s. Palaeogeography, Palaeoclimatology, Palaeoecology 64:69-78. 

Jones, P.D., K.R. Briffa and T.M.L. Wigley. 1988. Climate over Europe during the summer of 
1816. In: The year without a summer? Climate in 1816, An International Meeting 
Sponsored by the National Museum of Natural Sciences, Ottawa, 1988. Abstracts. p. 34. 

Jozsa, L.A., E. Oguss, P.A. Bramhall and S.G.Johnson. 1983. Studies based on tree ring data. 
Report to Canadian Forestry Service, Forintek Canada Corporation, 33 pp. 

LaMarche, V.C. Jr. and K. Hirschboeck. 1984. Frost rings in trees as records of major volcanic 
eruptions. Nature 307:121-126. 

Lamb, H.H. 1970. Volcanic dust in the atmosphere: with a chronology and assessment of its 
meteorological significance. Philosophical Transactions of the Royal Society of London 

Luckman, B.H. 1982. Little Ice Age and oxygen isotope studies in the Middle Canadian Rockies. 
Report to Parks Canada, Ottawa. 31 pp. 

. 1988. Dating the moraines and recession of Athabasca and Dome glaciers, Alberta. Arctic 
and Alpine Research 20:40-54. 


Luckman, B.H., J.P. Hamilton, L.A. Jozsa and J. Gray. 1985. Proxy climatic data from tree 

rings at Lake Louise, Alberta: a preliminary report. Geographie physique et Quaternaire 

Luckman, B.H., L.A. Jozsa and P.J. Murphy. 1984. Living seven-hundred-year-old Picea 
Engelmannii and Pinus albicaulis in the Canadian Rockies. Arctic and Alpine Research 

McCarthy, D.P. 1985. Dating Holocene Geomorphic Activity of Selected Landforms in the 
Geikie Creek valley, Mount Robson Provincial Park. M.Sc. thesis, University of Western 
Ontario, London, Ontario. 304 pp. 

Parker, M.L. 1985. Investigating the possibility of a relationship between volcanic eruptions and 
tree growth in Canada. In: Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 

Parker, M.L. and W.E.S. Henoch. 1971. The use of Engelmann spruce latewood density for 
dendrochronological purposes. Canadian Journal of Forest Research 1:90-98. 

Parker, M.L., L.A. Jozsa, S.G. Johnson and P.A. Bramhall. 1981. Dendrochronological studies 
of the coasts of James Bay and Hudson Bay. Jn: Climatic Change in Canada 2. C.R. 
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B.A. thesis, Geography Department, University of Western Ontario, London, Ontario. 
147 pp. 


How Did Treeline White Spruce at Churchill, Manitoba Respond to 
Conditions around 1816? 

David C.F. Fayle', Catherine V. Bentley’ and Peter A. Scott’ 


Annual radial increment throughout the stem, and height increment of individual white spruce 
trees at Churchill, Manitoba were reconstructed through measurement of ring widths on sections 
taken at close intervals throughout the stem. For the period around the eruption of Tambora in 
1815, four trees each from open-forest and forest-tundra sites provided data. On each site, one 
tree was less than 1 m in height in 1816, the others ranging from 3 to 6 m. Growth of the larger 
trees, as indicated by height and radial increment, was generally declining over the two decades 
prior to 1816. In the upper stem, particularly of the forest-tundra trees, radial increment was least 
in 1818. Effects were less severe in the lower stem and recovery in open-forest trees had begun 
in 1818 after a low in 1817. Net-height gain of the forest-tundra trees during 1816-20 was 
one-third that of the previous five years, whereas in open forest trees it more than tripled relative 
to reduced growth in the previous five years. In combination with the radial-increment data, this 
suggests the occurrence of conditions in 1816, or possibly late summer of 1815, that led to 
damage of the terminal bud and upper crown with loss of foliage and (or) reduction of foliar 
efficiency and production of new foliage. Such effects were much less severe on open-forest 
trees. The decline in overall tree growth was statistically significant in 1817-18 compared with 
the variability in tree growth for 10 years prior to 1815. Comparisons made with the period 
around 1835 (eruption of Coseguina) show subsequent growth reductions were greater than after 


The relationship between climatic variability and tree-ring widths is often difficult to establish and 
unclear at best. However, this relationship can be somewhat clarified by sampling 
climate-sensitive trees found in treeline areas where the annual energy deficit is restricting to 
growth (e.g., Jacoby and Ulan 1981); inclusion of other tree growth parameters may add 
considerable information from the outset (for a review see Fritts 1976). 

A severe climatic anomaly that coincides with a large volcanic explosion, such as reported for 
the eruption of Tambora during 1815 (Rampino and Self 1982; Catchpole 1985; Parker 1985; 
Wilson 1985), offers an opportunity to identify anomalous patterns of tree growth that coincide 
with the event (Parker 1985; Filion et al. 1986; Lough and Fritts 1987). By examining 
representative samples of tree populations, damage to the forest can be assessed which not only 
indicates the climatic impact of such an eruption, but also reveals information on how the climatic 
conditions may influence the forest environment. 

' Faculty of Forestry, University of Toronto, Toronto, Ontario MSS 1A1, Canada. 
2 R.R. 1, P.O. Box 22, Churchill, Ontario LOL 1K0, Canada. 
3 Department of Zoology, University of Toronto, Toronto, Ontario MSS 1A1, Canada. 



The field methods and development of the subsequent tree-growth index have been documented 
elsewhere (Scott et al. 1988). Briefly, 11 white spruce [(Picea glauca) (Moench) Voss)] and two 
tamarack [Larix laricina (Du Roi) K. Koch] were harvested in 1982 from five sites near the 
treeline at Churchill, Manitoba (58°45’N, 94°04’W). The trees were all open grown and ranged 
in age from 88- to 347-years old near their bases. Where identifiable in the upper stem, the 
lengths of annual height increments were measured and cross sections cut from their mid-point. 
Elsewhere, sections were cut at 10-cm intervals throughout the stem, except where branches were 
present. The sections were air dried, sanded and the ring widths measured to 0.01 mm on the 
four cardinal directions with a Holman DIGIMIC (Fayle et al. 1983). 

Specific volume increment (SVI) was a measure of the metabolic activity for a tree in 
each year (Shea and Armson 1972). This is the annual volume of wood produced relative to the 
surface area of the cambium that produced it (Duff and Nolan 1957); mathematically SVI is the 
average width of the growth layer. An advantage of SVI is that it is not a unidimensional 
parameter, such as ring width, because it integrates both diameter and height. Furthermore, since 
the reference point is a unit area of cambium, a common base is provided for comparison 
between trees. 

To develop the tree-growth index, the SVI series for each tree was standardized with a robust 
estimator (Draper and Smith 1981) using a negative exponential or, in the case of negative 
indices, a straight line of negative slope or through the average. The standardized SVIs were 
converted to ratios of the individual growth curves and then averaged to produce the final growth 

The final growth index was based on all trees sampled. However, only four of the white spruce 
from the open forest and four from the forest-tundra were present around 1816. Three from each 
type were greater than 3 m in height at that time and were used to analyze radial-longitudinal 
patterns of increment in relation to possible influences of climate. The fourth tree from each type 
was less than 1 m in height around 1816, and did not provide sufficient information for this 
particular purpose. 

Results and Discussion 

All of the trees show a decline in SVI of varying magnitude either during 1815 or in 1816 which 
persists for one to three years following (Figure 1, top and centre). The individual lags in 
response and magnitude do not allow for immediate conclusions regarding conditions during the 
summer of 1816. However, if we examine the overall status of the regional tree-growth index 
10 years prior to 1815, the growth during 1817 and 1818 is below the 95% confidence interval 
from what would be expected (Figure 1, bottom). The inference that a volcanic eruption may 
influence tree growth is strengthened by repeating the confidence interval test for the 1835 
period. The eruption during 1835 of Coseguina, which is much closer to Churchill than Tambora, 
may have had more potential for a stronger impact. In fact the 1835 period is the largest sudden 
decline in growth throughout the 1710-1982 period of the index. 

The cumulative net-height growth patterns for the individual open-forest trees do not indicate any 

consistent deleterious effect subsequent to 1815 (Figure 2a). Indeed, height increment appeared 
to be slowing down during the previous decade and recovered shortly thereafter (Figure 2b). In 


contrast, net-height growth was affected in the forest-tundra trees, where it was reduced for 
several years before recovering in the 1820s. The greater loss of terminal growth on the 
forest-tundra than on the open forest trees is reinforced by the similarity in pattern after 1835 
(Figure 2b); a substantial net gain in height did not occur on the forest-tundra trees for two 

The yearly longitudinal distribution of ring width throughout the tree stems for 1815-20 shows 
that changes did not occur uniformly (Figure 3). Reductions from 1815 to 1818 were greatest in 
the upper rather than lower stem, and more severe on the forest-tundra than on open-forest trees. 
The occurrence in the upper part of the growth layers of a ‘bulge’ in ring width, such as in 1819 
for Al and W2, may be the influence of lateral-branch development and (or) an increase in foliar 
amounts following damage to the current terminal or existing foliage. 

Minimum widths throughout the stem of the average forest-tundra tree occurred in 1818 with a 
58%, 50% and 27% reduction compared to 1815 in the upper 0.5 m, upper 0.5-2.0 m, and basal 
0.5-2.0 m respectively (Figure 4). In the average open-forest tree, the minimum occurred in 1818 
in the upper stem, but recovery was underway in the lower stem with the minimum occurring in 
1817; reductions during 1815-18 were 24%, 14% and 3% for the upper 0.5 m, upper 0.5-2.0 
m and basal O.5-2.0 m respectively. 

The reductions in height growth and in ring width in the upper stem subsequent to 1815 and 1835 
(Figure 3) indicate damage to, or loss of, the terminal buds and of foliage. The contribution of 
photosynthates and growth hormones by a branch to stem growth is related to the amount and 
proportion (by age) of the foliage it bears, the distance of this foliage from the stem and the 
amount of light it receives. A relatively short branch system with a high proportion of well-lit 
young foliage will make a high contribution to stem growth. 

In white spruce, the number of new needles that will be produced in the current year, and the 
potential shoot elongation were determined when the bud was formed in the late summer of the 
previous year (Owens et al. 1977). The degree of elongation and production of photosynthate are 
determined by conditions in the current year. Needles can be retained for 10-15 years at Churchill 
but there is a loss of photosynthetic efficiency with age. Current, one- and two-year- old foliage 
may contribute as much as 60% of the total assimilation in white spruce (Clark 1961). The 
influence of a favourable or unfavourable part of or whole growing season will therefore not only 
have different, but also lag effects, on growth, which can be compounded if there is a physical 
loss of new needles or premature loss of old needles. 

We have found that the loss of the terminal bud or shoot in treeline white spruce at Churchill is 
a common phenomenon often occurring at the same time as reduced width of the growth layer 
throughout the tree, but particularly in the upper stem. Poor growth occurs for several years 
while a lateral bud or branch establishes itself as the new terminal. Loss can come about through 
direct or indirect causes. An example of the latter could occur if the roots remain frozen while 
growth is under way (e.g., Scott et al. 1987), leading to desiccation and death of the needles and 
buds in the entire upper part of the tree (Sakai 1970; Kullman 1988). The proportion affected will 
determine the degree of growth reduction and, in combination with ongoing climatic conditions, 
the length of time to full recovery. The difference between open-forest and forest-tundra trees 
may be in the more exposed nature of the latter and the longer retention of older needles on the 
former, which provides a greater reserve. It is clear, from the slow recovery after the decline in 
growth during 1835, that many trees were apparently damaged this way, although there is little 
evidence of this occurring subsequent to 1815. 



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Figure 1: The specific volume increments (SVI) of the four open-forest and forest-tundra white spruce 
(top) that compose the growth index during the period around 1816. (This index, shown in the 
centre, is based on the SVIs of 13 trees.) Enlargement of the 1805-45 years (bottom) includes 
the mean (dashed line) and upper and lower (dotted line) 95 % confidence limits for the 10-year 
periods prior to the 1815 eruption of Tambora and the 1835 eruption of Coseguina, to show that 
the years following these eruptions exhibit unusually poor growth. 


SVI (cm?.cm?) 












W2 Figure 2a: Cumulative-height 






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190525505550 50504 
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curves for the three 
open-forest (top) and 
forest-tundra (bottom) 
trees and their 
respective averages 
(heavy line) during 
1800-30. The numbers 
at the left give the 
total height of the 
trees in 1800, to the 
nearest 3 cm. The 
circles and squares 
indicate section 
heights from which 
the curves were 
constructed. The 
vertical line marks 




1801-5 6-10 11-15 16-20 21-25 26-30 31-35 36-40 41-45 46-50 51-55 56-60 61-65 66-70 


Figure 2b: The net-height increase during five-year periods for 1801-70 inclusive for open-forest (hatched 
bar) and forest-tundra (open bar) trees. The five-year periods immediately following the 

eruptions of Tambora and Coseguina are underlined. 









1835 36 37 







= i.e) ow py ua 
_——— a 
pee ene 7 
[ee ef, 

1835 36 37 38 39 40 

Figure 3: The width of the growth layer (average of four radii) for the three open-forest (OF) and three 
forest-tundra (FT) trees during 1815-20, and for one example of each for 1835-40. 


: of TOP o-socm 

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: Nee TOP 50-200 cm 



BASAL 50-200CmM 

1800 10 20 30 

Figure 4: The average ring width for the 0-0.5 m and 0.5-2.0 m intervals from the contemporary apex, 
and for the 0.5 - 2.0 m interval above the stem base, for open-forest (dashed line) and 

forest-tundra (dotted line) trees. The vertical line marks 1815. 


The fact that radial growth is least in 1818 rather than 1816 is probably the result of cumulative 
effects arising from adverse conditions in 1816 or late summer of 1815. We suggest that the 
occurrence of conditions at that time led to damage of the terminal bud and upper crown with loss 
of foliage and (or) overall reduction of foliar efficiency and production of new foliage. Adverse 
conditions in August, after height increment was completed in either or both years, would reduce 
the amount of foliage produced the following year. Recovery of the open-forest trees below the 
apical 0.5 m in 1817, and a lesser reduction from 1817 to 1818 than from 1816 to 1817 in the 
forest-tundra trees, suggests that growing conditions were improving in late 1817. The reduced 
growth of the forest-tundra trees in 1818 may have been due to the cumulative effects of reduced 
foliar area, particularly in younger age classes, rather than adverse growing conditions per se. 
All trees showed improved radial growth in 1819 suggesting favourable conditions for bud 
development existed in 1818 (Figure 5). 


Figure 5: Diagrammatic presentation of reduction in shoot growth due to adverse effects on bud 
development, shoot and needle elongation without (left) and with (right) damage to the terminal 
bud, shoot and needles. Horizontal lines represent needles of different age classes, from current 
year (heavy line) to five-years old (dotted line). Older years are not shown. The brackets 
indicate the needle classes that would normally contribute the bulk of photosynthate. 

Year one (e.g., 1815) shows normal growth and bud development. In year two unfavourable 
conditions throughout the growing season restrict shoot and needle elongation and bud 
formation. Incipient damage to the bud and needles may occur. In year three, growing 
conditions are more normal and needle elongation and bud formation are not restricted, but the 
amount of new foliage is reduced due to previous adverse conditions. Where damage occurred, 
a lateral bud may begin to assume dominance, but its small size has restricted the amount of 
needles produced. In year four, growth and development are near normal. Where damage had 
occurred, the quantity of photosynthetically-efficient foliage is still low and ring width is 
minimal. Recovery occurs in year five here, whereas it was already underway in the undamaged 


The above scenario is complemented by the observations of Filion et al. (1986) who reported a 
high occurrence of ‘light rings’ in 1816 and 1817 in krumholz black spruce in northern Quebec. 
We have not had the opportunity yet to determine their presence in our trees. If they do occur, 
which is likely, unfavourable conditions in the late part of the growing season and (or) a shortage 
of photosynthate, the result for example of needle loss, are suggested. In the ‘light rings’ 
illustrated by Filion et al. (1986), the last formed tracheids in the annual ring show a normal, 
narrow radial diameter but wall thickening is minimal. A supply of photosynthate is required to 
complete the process of wall thickening and environmental conditions must permit the completion 
of the maturation process for normal latewood formation. 

From the examination of tree growth during 1815 and particularly 1835, it appears that a 
stochastic event, such as a volcanic eruption, occurring many thousands of kilometres away may 
have a widespread detrimental effect on forest productivity. Climatic conditions at Churchill 
following the eruption of Tambora in 1815 and Coseguina in 1835 did have adverse effects on 
growth of white spruce. Correspondingly, Wilson (1985) reports that, on the east side of Hudson 
Bay, conditions were poor during the late summer of 1815, during 1816 and possibly the first 
half of 1817. Similarly, while it is apparent that 1816 was not truly without a summer at 
Churchill, it may have been one of 5°C temperatures instead of the long-term average of 10°C. 


We thank Ed Cook and Gordon Jacoby for helpful advice and supplying some analysis programs. 
Ring-width measurements were made using facilities of the Ontario Tree Improvement and Forest 
Biomass Institute, Ontario Ministry of Natural Resources, with grants supplied to the authors by 
Environment Canada and to P. Scott by Indian and Northern Affairs Canada. NSERC provided 
travel funds for D. Fayle. We also thank Roger Hansell for his help in the project and C.R. 
Harington for his support. 


Catchpole, A.J.W. 1985. Evidence from Hudson Bay region of severe cold in the summer 
of 1816. In: Climatic Change in Canada 5. C.R. Harington (ed.). Syllogeus 55:121-146. 

Clark, J. 1961. Photosynthesis and respiration in white spruce and balsam fir. Syracuse 
University, State University College of Forestry Technical Publication 85:1-72. 

Draper, N. and H. Smith. 1981. Applied Regression Analysis. Second Edition. John Wiley 
& Sons, Inc., Toronto. 709 pp. 

Duff, G.H. and N.J. Nolan. 1957. Growth and morphogenesis in the Canadian forest 
species. II. Specific increments and their relation to the quantity and activity of growth in 

Pinus resinosa Ait. Canadian Journal of Botany 35:527-572. 

Fayle, D.C.F., D.C. MacIver and C.V. Bentley. 1983. Computer-graphing of annual ring 
widths during measurement. The Forestry Chronicle 59:291-293. 

Filion, L., S. Payette, L. Gauthier and Y. Boutin. 1986. Light rings in subarctic conifers as a 
dendrochronological tool. Quaternary Research 26:272-279. 


Fritts, H.C. 1976. Tree Rings and Climate. Academic Press Inc. New York, New York. 
567 pp. 

Jacoby, G.C. and L.D. Ulan. 1981. Review of dendroclimatology in the forest-tundra 
ecotone in Alaska and Canada. In: Climatic Change in Canada 2. C.R. Harington (ed.). 
Syllogeus 33:97-128. 

Kullman, L. 1988. Subalpine Picea abies decline in the Swedish Scandes. Mountain 
Research and Development 8:33-42. 

Lough, J.M. and H.C. Fritts. 1987. An assessment of the possible effects of volcanic 
eruptions on North American climate using tree-ring data, 1602 to 1900 A.D. Climatic 
Change 10:219-239. 

Owens, J.N., M. Molder and H. Langer. 1977. Bud development in Picea glauca. I. 
Annual growth cycle of vegetative buds and shoot elongation as they relate to date and 
temperature sums. Canadian Journal of Botany 55:2728-2745. 

Parker, M.L. 1985. Investigating the possibility of a relationship between volcanic eruptions 
and tree growth in Canada (1800-1899). In: Climatic Change in Canada 5. 
C.R. Harington (ed.). Syllogeus 55:249-264. 

Rampino, M.R. and S. Self. 1982. Historic eruptions of Tambora (1815), Krakatau (1883), 
and Agung (1963), their stratospheric aerosols, and climatic impact. Quaternary Research 

Sakai, A. 1970. Mechanism of desiccation damage of conifers wintering in soil-frozen areas. 
Ecology 51:657-664. 

Scott, P.A., C.V. Bentley, D.C.F. Fayle and R.I.C. Hansell. 1987. Crown forms and 
shoot elongation of white spruce at the treeline, Churchill, Manitoba, Canada. Arctic and 
Alpine Research 19:175-186. 

Scott, P.A., D.C.F. Fayle, C.V. Bentley and R.I.C. Hamsell. 1988. Large scale changes in 
atmospheric circulation interpreted from patterns of tree growth at Churchill, Manitoba, 
Canada. Arctic and Alpine Research 20:199-211. 

Shea, S.R. and K.A. Armson. 1972. Stem analysis of jack pine (Pinus banksiana Lamb.): 
techniques and concepts. Canadian Journal of Forest Research 2:392-406. 

Wilson, C. 1985. The Little Ice Age on Eastern Hudson/James Bay: the summer weather 
and climate at Great Whale, Fort George and Eastmain, 1814 to 1821, as derived from 
Hudson’s Bay Company records. Jn: Climatic Change in Canada 5. C.R. Harington (ed.). 
Syllogeus 55:147-190. 


The Climate of Central Canada and Southwestern Europe Reconstructed 
by Combining Various Types of Proxy Data: a Detailed Analysis of the 
1810-20 Period 

J. Guiot' 


In this study, an attempt is made to synthesize various kind of proxy records and to reconstruct 
complete climatic series. In central Canada, tree-ring series and historical records from the 
Hudson’s Bay Company (i.e., ice-condition and early instrumental data) have been assembled to 
reconstruct a seasonal temperature and sea-level pressure network back to 1700. The summer of 
1816 was among the coldest in the period studied. The beginning of the nineteenth century was 
also cold, especially after 1807, but the main characteristic is great variability (e.g., 1818 was 
one of the warmest years since 1700). These results are compared with those obtained in a similar 
manner for southwestern Europe and northwestern Africa on the basis of tree-ring series, '*0 
records, wine-harvest and other archival data. The 1810-20 period was also among the coldest 
of the last millennium, and 1816 was one of the four coldest years since the eleventh century. In 
the Mediterranean region, this period was far less cold. Some details are given on the method 
used for these reconstructions. As the proxy series are not homogeneous, particular devices are 
needed to estimate missing data and to reconstruct low-frequency components. The techniques 
are adapted from multiple regression, digital filtering, bootstrap analysis and principal component 

The Data in Central Canada 

The meteorological network is made up of 67 stations selected from the meteorological database 
of the Atmospheric Environment Service of Canada. The period of analysis is restricted to 
1925-83 and the region is delimited by 61° - 105°W and 47° - 73°N. The monthly data are 
averaged into seasonal series. 

The second Canadian data set is built from the proxy series (Figure 1) available at the date of the 
study (Guiot 1985a), including: 

e freeze-up and break-up dates of rivers entering the western shore of Hudson and James bays: 
nine series (1714-1871 at maximum) derived by Catchpole and Moodie (1975) and extended 
to the modern period using recent data (Allen 1977); 

e freeze-up and break-up dates of the Red River at Winnipeg, extending from 1798 to 1981: two 
series build by Rannie (1983); 

¢ monthly temperature data for York Factory (1774-1910) and Churchill Factory (1768-69/ 

1811-58) derived by Ball and Kingsley (1984), spatially and temporally averaged into four 
seasonal series for the York-Churchill region. 

' CNRS UA 1152, Laboratoire de Botanique Historique et Palynologie, Faculté de St. Jér6me, 13397 Marseille 
Cedex 13, France. 


These data are completed by tree-ring series. The trends of these series are modeled by negative 
exponentials, polynomials or filtered curves by the authors as proposed by Fritts (1976). Indexed 
series are obtained by dividing each ring width by its trend. The series used are the following: 

© two white spruce ring-width indices series from Nain, Labrador (1769-1973) and Border 
Beacon, Labrador (1660-1976) from Cropper and Fritts (1981); 

¢ one larch ring-width indices series from Fort Chimo (now Kuujjuaq), Québec (1650-1974) also 
from Cropper and Fritts (1981); 

© two white spruce series from Cri Lake, near Kuujjuarapik, Québec, (1750-1979) (Parker er 
al. 1981), the first being ring-width indices and the second being ring maximum densities; 

¢ two white spruce ring-width indices from Churchill, Manitoba (1691-1982) by P. Scott in 
Hansell (1984), the first was sampled in open forest and the second in forest-tundra. 

Finally a total of 22 proxy series are available to reconstruct temperature in central Canada for 


Location of the sites 
A: Archives sites 

Wl: Tree-rings sites 

° ° 
== = E | 
< repel Island 2 

Figure 1: Location of the proxy-series sites in Canada. 

The Data in Europe and North Africa 
Meteorological data are the annual series gridded by Jones et al. (1985) extending from 35° to 

55°N by steps of 5°, and from 10°W to 20°E by steps of 10°. So 20 series are available from 
1851 to 1984, with missing data mainly before 1900. 


The second data set, the proxy series, are collected from the longest proxy series existing for 
Europe and Morocco (Figure 2). A part of them consists of tree-ring chronologies of various 
species from various sites. They are also detrended as suggested by Fritts (1976). The set 

oak ring-width series from west of the Rhine, near Trier, Germany (820 to 1964) collected 
and indexed by Hollstein (1965); 

oak ring-width series from the Spessart forest area (SO0°N, 9°30’E) in Germany (840 to 1949) 
collected and indexed by Huber and Giertz-Siebenlist (1969); 

oak ring-width series from Belfast, Northern Ireland (1001 to 1970) collected and indexed by 
Baillie (1977); 

oak ring-width series from southwestern Scotland (946 to 1975) collected and indexed by 
Pilcher and Baillie (1980); 

pine ring-width series from southern Italy (1148 to 1974) collected and indexed by 
Serre-Bachet (1985); 

larch ring-width series from Vallée des Merveilles, southern French Alps (1100 to 1974) 
collected and indexed by Serre (1978); 

fir ring-width series from Mont Ventoux, southern France (1660 to 1975) collected and 
indexed by Serre-Bachet (1986); 

pine ring-width series from northern Italy (925 to 1984) collected and indexed by Bebber 
(personal communication); 

larch ring-width series from Orgére, northern French Alps (1353 to 1973) collected and 
indexed by Tessier (1981); 

two larch ring-width series from Mercantour, southern French Alps (1701 to 1980 and 1732 
to 1981) collected by Guibal (personal communication) and indexed for this study. 

Another group of proxy series is composed of data derived from archives. These historical data 
have been compiled by various historians and/or climatologists: 

decadal temperature estimates of Bergthorsson (1969) for Iceland (1050 to 1550). These data 
were analyzed by Ogilvie (in Ingram et a/. 1978), and those before 1170 and after 1450 were 
reported as unreliable - the unreliable decades are considered as missing; 

summer temperature index of Bray (1982) based on German and French wine-harvest data and 
central England (Manley) temperatures (1453 to 1973); 

the Pfister (1981) thermal indices in Switzerland, averaged on an annual basis from 1550 to 


@ Three glaciological series (18O) 
® @ Greenland 

®@ Iceland 

Sources of Data 
a Tree-rings 
m Grape-harvest dates 

nA me 


- @ Other data 
x Temperatures at 
1 gridpoints 
* A x 
1 = \ 
i a's I he Grape-harvest series 
; ui! ||! (France, Rhine, 

} J” Switzerland) 

10oW "aes 20°E 
es 0° 10°E 
17 Tree-ring series 

Figure 2: Location of the proxy-series sites in Europe and Morocco. 
e the mean annual dates at the beginning of the grape harvest in northeastern France, French 
Switzerland, and southern Rhineland of Le Roy Ladurie and Baulant (1981) (1484 to 1879); 

e the average annual dates at the beginning of the grape harvest in Switzerland reported by 
Legrand (1979) (1502 to 1979); 

e frequency of southwesterly surface winds in England (1340-1978) from direct observations 
(1669 to 1978) in the London area and from historical proxy data before. These data are 
reconstructed by Lamb (1982); 

A last category is provided by '*O data in the Arctic ice. These isotopic series can be considered 
as good indicators of temperature, since the condensed vapour is enriched in heavy isotopes: 

¢ Camp Century, Greenland, '*O quasi-decadal values (1200 to 1970) collected and analyzed by 
Dansgaard et al. (1971); 

© two isotopic series in central Greenland, 30-year running means of annual maxima of '*O in 
ice cores compiled by Williams and Wigley (1983) (1180 to 1800). 


Finally, to these data are added the first three principal components of the 17 longest cedar 
ring-width series in Morocco, sampled by A. Munaut and C. Till and analyzed by Till (1985). 
These series (1068 to 1979) represent nearly 40% of the total variance of the 17 raw series. 

The period 1068-1979 is retained for a total of 23 series. To simplify matters, European data will 
include both European and northwestern African data. 

Data Conditioning 

The predictand matrix as well as the predictor contains missing data. Therefore it is fairly natural 
to estimate the gaps before beginning any detailed analysis. For the Canadian data, the method 
employed to estimate the missing data is explained in Guiot (1985a) and, with more details, in 
Guiot (1986). The general procedure is similar to that used for the management of the European 
series, described here. 

The best analogues method is used to estimate the missing data of the proxy-series matrix. The 
main advantage of this method is that we have not to assume any linear relationship between the 
variables. This is particularly recommended when, like here, the series are highly heterogeneous. 
The estimate of a missing observation for a given series is provided by the most similar 
observations (analogues) of the same series within the 1200-1900 interval, the distance between 
observations being established on the m,, observations of the series available. 


d= Lo, -x,? (1) 

The observations, denoted by k, available among the 20 best-fit analogues of observation i 
provide the wanted estimate 

Dk x, fd? 
a= plea Se (2) 

Yk d2 

The correlation between estimates and actual values computed on the available data and averaged 
on the 23 series is 0.73 (ranging from 0.45 to 0.86), which is highly significant. For observations 
outside the 1200-1900 period, the mean correlation remains high, say 0.60 (ranging from 0.22 
to 0.89). It must be noted that we have not estimated any coefficients so that the statistics 
computed on the calibration interval as well as on the verification one can be considered as 
independent. The mean and the standard deviations of the estimates are quite close to the actual 
ones, with discrepancies less than 15% of the mean standard deviation. Depending on the number 
of degrees of freedom, we can consider that the estimates are reliable. 

For the meteorological data matrix, multiple regression was used. This method cannot be applied 
directly because the number of regressors is not constant on the total calibration interval (1851- 
1984). If m, is the number of regressors available for observation i (i.e., with no missing data), 
the regression equation may be written as follows: 


=a, + yy By ee (3) 
k= 1 

The correlation between estimates and actual values averages 0.76, with the highest values in the 
Northwest (0.90). These coefficients are highly significant, but the estimates must be considered 
as less reliable at the southern margin of the region analyzed (correlation around 0.70). 

Extrapolation of the Temperature Series 

When the predictor and predictand matrices, are fully determined, it is possible to extrapolate the — 
annual-temperature series from the proxy series, using the common observations to calibrate a 
relationship. It is advisable first to transform the raw series into principal components (PCs). 
Indeed, a large proportion of the high order PCs represents extremely small proportions of 
variance, so that they can be assumed to be indistinguishable from statistical noise. 

Reduction of the Number of Variables 

For the European annual-temperature series, 10 PCs are used explaining together around 90% 
of the variance. For the European proxy series, 19 PCs are used explaining 95% of the variance. 
For the Canadian season temperature series, the first four PCs used explain between 82.5% (for 
summer) and 91% (for autumn) of the total variance. For the proxy series, the number of PCs 
depends on the season reconstructed. 

Bootstrap Regression 

In central Canada, a multiple regression has been employed to calibrate the relationship between 
climate and proxy series. In Europe, a more sophisticated approach, termed bootstrap regression, 
seemed more advantageous. 

Bootstrapping is a recent technique devised by Efron (1979) to estimate statistics for unknown 
population distributions by Monte Carlo simulations. The idea is to resample the original 
observations in a suitable way to construct pseudo-data sets on which the estimates are made. In 
regression, this is particularly useful when the residuals are non-normal or autocorrelated, or 
when the data set is too small. 

The bootstrap method is in fact a generalization of jackknife replication. The frame of the method 
can be summarized in a few lines. From the interval (1,n), where n is the size of the original data 
set, n pseudorandom numbers are randomly taken with replacement using a uniform distribution 
protocol. These n numbers are used to resample the actual observations. We should insist here 
on the fact that an observation is the vector of the m proxy data and p climatic parameters 
corresponding to the same year. The n observations selected in this way provide a pseudo-data 
set. This is repeated an arbitrary number NC times, and at each time, a regression is computed. 

The reliability of a particular statistical model must be assessed by calculating a number of 
verification statistics measuring the degree of similarity between predictand observations and their 
estimates for time periods independent of the calibration. So a successful reconstruction is one 
for which it is demonstrated that independent estimates continue to be accurate at a level greater 
than would be expected solely by chance. "The process used to optimize the coefficients of the 
model virtually ensures that the results will be more accurate for the calibration data than for any 
other observations to which it may be applied. It is why the decreasing of accuracy should be 


measured whenever possible.” (Fritts and Guiot 1988). Bootstrap regression enables one to 
integrate verification in the calibration process and to use the n observations both for the 
calibration and the independent verification: 

e for each of the NC replications, the regression coefficients are computed and applied to proxy 
series to obtain the corresponding reconstruction; 

e the reconstruction is compared to the actual climatic series both on the set of retained 
observations and on the others; thus verification statistics are calculated NC times; 

e the mean and standard deviations of the verification statistics are obtained on the dependent 
and independent data set; 

e the final reconstruction is given by the median of the NC replicated reconstructions, and a 
90%-confidence interval is given by the 5th and 95th percentile. 

Decomposition of the Spectra into Two Bands 

Before computing a bootstrap regression, the predictors and the predictands are filtered, so that 
their spectra are decomposed into two bands (Guiot 1985b). Once more, the method is illustrated 
with European data. We use a nine-weights low-pass filter with a cut-off period of seven years. 
The effect of this filter is illustrated in Figure 3 with the first PC of the proxy series. The 
complementary high-pass filter enables us to retain the short-term fluctuations of the series. The 
raw Series is the sum of both low-frequency and high-frequency components (Figure 3). 

In the two frequency bands, bootstrap regressions are calibrated on the common period, 
1851-1979. This method is particularly necessary for the low-frequency components dominated 
by large autocorrelations, which induce troubles in the interpretation of the fit quality. The 
"abnormality" of these smoothed data is compensated for by a lot of simulations. 

Table 1 presents some Statistics useful for the evaluation of the regressions. For each of the 50 
simulations, the estimated means and standard deviations are compared to the actual ones on the 
randomly-drawn observations, as well as on the others. The deviations of these statistics are 
averaged over the 50 simulations (Table 1). Apparently the standard deviations are slightly 
underestimated, as expected, and the biases are not greater on the independent observations. 
Concerning the calibration data set, the correlations between estimated and actual observations 
are lower for the high-frequency components than for the low-frequency components. 
Nevertheless this must be appreciated regarding the reduced number of degrees of freedom of 
autocorrelated series. The most important feature is the lack of stability, appearing in the 
independent data set, of the high frequencies for components 3 to 6 and 8 to 10, while the low- 
frequency components estimates are quite stable. This justifies the spectral decomposition. 

The regression coefficients are applied in each band to extrapolate the 10 temperature PC series 
back to 1068. The entire spectra are recomposed by adding the reconstructed high-frequency PC 
series to the low-frequency ones. Table 2 presents the effect of this addition for two periods in 
the calibration period. During the first period (1851-1900), temperature observations are less 
abundant and of lesser quality than during the second (1901-79). The means and standard 
deviations in the older period appear to be systematically more underestimated than in the more 
recent one by a factor of 2. These underestimates are negligible for the means: they represent 
between 0.7 and 2.5% of the total variance of the PCs (that is 2.10*). The biases are higher for 


1380 130 Nay ma 

RA ic 1 

‘Ll FREQ. Foal 

“HIGH FREQ. PC t" sane uneeailonaniecet 3 sega ota de shed ea 3 ere hee 

Figure 3: The first principal component of the European proxy series (1850-1979) and its spectral 


Table 1: Verification Statistics for the Reconstruction of the Low- and High-Frequencies Component 
of the First 10 PCs of the Annual Temperatures. [These statistics are averaged on the 50 
replications: (a) on the calibration observations (randomly drawn); (b) on the others. dM = 
estimated mean minus actual mean; dS = estimated standard deviation minus actual standard 
deviation; R = correlation coefficient between estimates and actual variables + 1 standard 
deviation. | 

Low Frequencies 

Var. 1 2 3 4 5 6 7 8 9 10 
dM (a) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 
(b) -0.01 0.01 -0.01 0.00 0.00 0.04 0.01 0.01 -0.01 -0.01 

dD (a) 0.15 -0.28 -0.30 -0.13 -0.33 -0.21 -O0.17 -0.22 -0.34 -0.17 
(6) -0.15 -0.30 -0.29 -0.11 -0.33 -0.22 -0.17 -0.22 -0.33 -0.19 

R (a) 0285), (0572) 10:70) (058740167 0379 0:83 OL78 0:66 (0:83 
+ 0.03 0.03 0.03 0.02 0.04 0.03 0.03 0.03 0.04 0.03 

(6) (0574) (0553: (0/57 ~20:83 0.48 «420165 0.75 0.67 0:47 0.75 

+ 0.09 0.09 0.09 0.04 0.10 0.09 0.05 0.07 0.09 0.07 

High Frequencies 

Var. 1 2 5) 4 =) 6 U 8 9 10 
dM (a) 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 
(b) 0.02 -0.01 0.04 -0.01 0.01 0.03 0.00 -0.01 -0.01 0.03 

dD (a) -0.35 -0.50 -0.44 -0.54 -0.49 -0.62 -0.42 -0.48 -0.57 -0.51 
(b) -0.29 -0.45 -0.41 -0.54 +-0.43 -0.56 -0.40 -0.43 -0.53 -0.49 

R @)) 1G:65" 0:50 “0:56 0:46-0:51) 0:38 0:58 0.52 0:43 0:49 
+ 0.04 0.05 0.05 0.06 0.04 0.06 0.04 0.04 0.07 0.05 

(G) 0F48 «©0224 0:18 0.127 0.19 -0:10 0:31 OS -0:01 O71 
30309). 10:09) Onis FO:15* 0:10) OF1t, 0:09 ~=—0209) Ot. OF 11 

the variability: between 10 and 20% of the total variance. These results are clearly better than 
those obtained without spectral decomposition. With this last method, the correlation gain is about 
0.14, and the underestimating factor is divided by 2. We conclude that spectral decomposition 
increases the quality of fit, but we cannot infer that this best fit is warranted on independent 
periods. The example dealt with in Fritts and Guiot (1988) nevertheless confirms the stability of 
such extrapolations. 

The reconstruction of the raw-temperature series is obtained by postmultiplying the 10 PC series 
matrix by the eigenvector matrix. The mean correlation between estimated and actual values is 
0.63 (from 0.46 to 0.76) on the 1851-1979 period, with a maximum of more than 0.70 in the 
Northwest. These reconstructed series are provided with a 0.90-level confidence interval that is 
0.3°C in mean. 

Analysis of the Reconstructions as a Whole 

The reconstructed series in Europe are analyzed from different points of view. The trend of the 
16 series, at latitudes ranging from 40° to 55°N is plotted in Figure 4, on the basis of 20-year 


Table 2: Verification of the Reconstructions of the 20 Annual-Temperature Series PCs (Multiplied 
by 100). The sum of the low and high frequencies are first verified, and then the regression, 
without separating low from high frequencies: (a) actual means; (b) estimated mean; (c) 
actual standard deviations; (d) estimated standard deviations; (e) correlation between actual 
and reconstructed PCs. SSD = sum of squared differences [sign (-) means underestimates]. 

Low and High Frequencies 

1901-1979 1851-1900 
a b ( d e a b c d e 

PCOl1 -41 -41 270 187 0.69 87 3i7/ 296 195 0.73 
PC02 33 8 167 95 0.61 -55 -31 170 89 0.51 
PCO03 -17 -19 154 88 0.50 12 11 190 85 0.61 
PC04 -15 -18 100 84 0.83 AH 20 104 41 0.57 
PCOS 13 -11 69 48 0.37 -24 6 . 148 75 0.79 
PC06 -34 -28 67 43 0.47 69 34 74 46 0.59 
PC07 2 -2 58 42 0.68 -3 -1 69 41 0.74 
PCO8 6 6 62 41 0.62 -16 -17 82 43 0.71 
PCO9 -7 10 70 38 0.58 -3 -6 80 29 0.47 
PC10 -8 -5 49 31 0.56 8 3 66 38 OW2 
SSD (-)1504 (-)20867 (-)5290 (-)43559 

mean 0.59 0.64 

Standard Regression 

1901-1979 1851-1900 
a b c d e a b c d S 

PCO1 -41 1 270 144 61 87 20 296 160 .65 
PC02 33 -1 167 33 535) -55 -1 170 40 23 
PCO3 -17 -21 154 67 37 12 19 190 oy 58 
PC04 -15 -7 100 66 .74 27 14 104 5 45 
PCOS is" -5 69 31 Oi -24 5 148 5h7/ .62 
PC06 -34 -2 67 26 33 69 18 74 36 41 
PC07 12 3 58 31 .62 -3 9 69 27 58 
PCO08 6 4 62 28 .47 -16 -14 82 26 51 
PCO9 -7 -3 70 21 38 -3 -9 80 19 .43 
PC10 -8 -1 49 18 Di 8 -4 66 18 .60 
SSD (-)4498 (-)53549 (-)11393 (-)83891 
mean .45 sey 

periods. A climate generally colder than now, results interrupted by a few warm periods. A 
spatial distribution of the anomalies is presented at some key 20-year periods for the 20 

Before 1200, the temperature was extremely low in the whole region. From 1200 to 1400, the 
analyzed region experienced a relatively warm climate mainly in the Southwest. This warm 
period is often called the "Little Climatic Optimum" of the Middle Ages. The results for the 
1070-1420 period are verified by comparing the indices of Alexandre (1987) that are valuable for 
Mediterranean and non-Mediterranean western Europe. In order to make these indices 
representative of both winter and summer, we have used the difference "Winter severity index 
minus Summer precipitation index" to represent the annual temperature. It confirms that a 


warming is obvious from the beginning of the thirteenth century, with a mean of -1.8 before 1220 
and 0.23 after (Figure 4). 

From 1420 to 1460, conditions were very cold, except in North Africa. The warming at the end 
of the Middle Ages lasted from 1460 to about 1550. The spatial distribution was similar to that 
of 1200-1400, with a maximum to the West. 

The period generally called the "Little Ice Age" seems to have begun about 1500. The first part 
of this period (1550-1610) was effectively cold mainly in areas along a diagonal extending from 
the British Isles to Tunisia. The seventeenth century was generally warm in the Southwest. In the 
Southeast, the cooling started around 1550. In fact, the Little Ice Age really started at the end 
of the seventeenth century. Temperature was low everywhere except in the Southwest. It lasted 
until 1860, with two particularly cold periods about 1700 and 1815. It had no equivalent in the 
Southwest, although the precipitation reconstructions of Till and Guiot (1988) indicate increasing 
moisture - particularly during these two extreme episodes. Richter (1988) confirms the climatic 
differences between the southwestern Mediterranean Basin and northwestern Europe. Its 
reconstruction of summer precipitation from pine tree-ring series shows that 1810-20 was wet in 
central Spain and its reconstruction of winter temperature shows that the same area was warm. 
As central Spain is located at the midpoint between Morocco and the rest of western Europe, 
these reconstructions are simultaneously, a confirmation of our temperature reconstructions and 
the precipitation reconstructions of Till and Guiot (1988). 

The modern warm period began in the mid-nineteenth century, with a maximum between 1930 
and 1950 - especially in the Northwest and in the Southeast. 

Similar reconstructions for the four seasons have been obtained in central Canada, but only for 
the last three centuries. The reconstructions are detailed in Guiot (1985a). WI focus here on a 
comparison with Europe. Figure 5 shows three synchronous long periods on both continents: 
1700-50; 1780-1820; and 1850-1920. After the beginning of the twentieth century, the general 
warming appeared in Canada some five years later than in Europe. The synchronism during the 
Little Ice Age could mean that it is forced by an external common phenomenon. 

The Year 1816 

If 1810-20 appears, as a whole, very cold in Europe, it is highly variable in Canada (Figure 5). 
For example, the summer was very cold in 1816 but it was very hot two years later. On the two 
continents, the cooling began at the beginning of the nineteenth century. 1816 is only a period 
where this cooling reached an extreme. If the volcanic eruption of Tambora in 1815 had an 
influence on the severe climate of the following year, it only accentuated a trend, and its eventual 
effects must be placed in the context of the cold period of the "Little Ice Age". This trend is 
noticeable as well in Canada as in Europe. Figure 5 also shows that the summer of 1816 was the 
coldest of the last three centuries in central Canada. In Europe, we have found three other years 
as cold as 1816: 1081, 1454 and 1703. The four coldest years of the millennium reached mean 
anomalies of -1.5°C (in the range 10°W - 20°E and 35° - 55°N). 

Central Canada reconstructions are sufficiently precise for a chronology of the cooling in the 
region to be established. Figure 6 presents the distribution of temperature anomalies in relation 
to 1950-79. The temperature of the North is a blank because the proxy series used is not 
representative of latitudes higher than 65°N. Winter 1816 was nearly normal in the whole region 
studied except in the Southwest where the cooling was already perceptible (negative anomalies 


(anomalies) hye. aceite 

EUROPE = ~-"-a2e@ 

—— oe ee ee oe 

2 Nite tall 

wee eees 
senor Pe 

sapliieens |I 

=e ee we Se 

. . . often * 
= Tee 
Pe ht) I i 

~" 43h) 


CeCe ee 


ween ne ne ee eg 

eee weer eens 
Ce ee 

Ce we er ey 
Pewee ere ee eee te Mm we eee 
secre eee ees Be vane 

-2 -1 O 1 WARM 
Alexandre Index 

Figure 4: The 20-year trend of temperature variations in Europe (area restricted to latitudes 40°-55°N 
and longitudes 10°W-20°E). The distribution of the anomalies for some characteristic periods 
is shown for the total area (including 35°N). The broken horizontal lines represent the 
90%-confidence intervals computed by bootstrapping. Between 1070 and 1410, are the 
smoothed Alexandre (1987) indices representing winter severity and summer precipitation. 








Figure 5; The annual temperature in southwestern Europe and central Canada (both being averages of the 
individual reconstructed series) and the summer temperature in central Canada. The series are 
smoothed with a digital filter (cut-off period = seven years). 




Figure 6: The distribution of temperature anomalies for the four seasons in central Canada. 

greater than 1°C). The cooling affected the whole region in spring, with a maximum in the 
Southwest where negative anomalies of 2°C are reached. In summer, the temperature anomalies 
were globally -1°C with a minimum of almost -3°C in the region of Kuujjuarapik (southeastern 
Hudson Bay). A secondary minimum of -2°C occurs in the Churchill region (western Hudson 
Bay). The Southwest has already begun to warm, since spring anomalies are -2°C and summer 
ones -1°C. Autumn is nearly normal everywhere except in the Southwest where the positive 
anomalies are +1°C. 

In Europe, because emphasis was laid on the ability to reconstruct temperature series Over a 
millennium, it is impossible to collect a sufficient number of series to obtain a seasonal 
resolution. Figure 7 is nevertheless instructive respecting the spatial differences of the cooling. 
Briffa et al. (1988; this volume) have already shown that the summer of 1816 in central Europe 
was less cold than in western Europe. This is confirmed as far as the annual temperatures are 
concerned, and it is possible to more precisely judge the temperature of the western 
Mediterranean Basin. The maximum negative anomalies concern the British Isles and northern 
France (-3°C), but they extend far away to Africa - especially Tunisia. This teleconnection 
between northwestern Europe and North Africa is a classical synoptic configuration, which is well 
known nowadays in southern France during "Mistral" and "Tramontane" winds. Northwesterly 
air masses are canalized under the influence of an anticyclone located over Spain and a low 
pressure centre over the southern Alps and Gulf of Genova. The winds accelerate and become 
drier down the Rhone Valley (Mistral) and by invading the area between Pyrénées and Massif 
Central mountains (Tramontane) so that their influence (when they are exceptionally strong) is 
sometimes felt in Corsica, and even in North Africa. Perhaps this meteorological situation 
occurred very often during the summer of 1816. However, the coldness of this year was not 
general: apparently Morocco and southern Spain were largely influenced by southerly winds since 
the negative anomalies are lower than 1°C. Central Europe, with a more continental climate, was 
also less affected by this general cooling. 


Figure 7: The distribution of the annual temperature anomalies in Europe and northwestern Africa in 


The summer of 1816 was the coldest in the last three hundred years in central Canada. The other 
seasons have been about normal or slightly colder. The greatest negative anomalies concern the 
area southeast of Hudson Bay, and the smallest ones the region north of Hudson Bay. In Europe, 
1816 was among the coldest years of the millennium with minimum temperatures (anomalies 
close to -3°C) extending from the British Isles to Tunisia. Northwesterly winds chilled western 
Europe: these cold air masses accelerated and dried between the Pyrénées and Massif Central 
mountains (Tramontane wind) and down the Rhone Valley (Mistral wind), crossing the 
Mediterranean Sea to Tunisia. At the same time, Morocco remained relatively warm. In central 
Canada, more details are available about 1816 from seasonal records. Apparently the cooling 
began in winter in the Southwest and ended by the close of summer, whereas it began a season 
later (in spring) in the East, also ending later (in autumn). 

The severe climate characteristic of this year must be placed in context. The cooling began a few 
years before 1816, at the beginning of the decade. Then aerosols from the volcanic eruption of 
Tambora only exacerbated a trend already existing in Europe and central Canada. Comparison 
of results from the two continents shows strong coherency in the low-frequency variations of 
temperature on both sides of the Atlantic Ocean during this globally-cold period of the Little Ice 
Age. The coherency is weak in warmer periods. 


This study shows how to synthesize various proxy series available to provide a better knowledge 
of past climatic changes in Europe. More records must be used in order to obtain maximum 
reliability of the gridded temperature reconstructions. It also appears that the Mediterranean 
climate, which is now very different from the northern European one, has been so for many 
centuries. Information concerning northern Europe cannot be directly extended to southern 
Europe. More proxy series related to the Mediterranean climate must be collected to achieve 
better reliability. 


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Climatic Conditions for the Period Surrounding the Tambora Signal in 
Ice Cores from the Canadian High Arctic Islands 

Bea Taylor Alt’, David A. Fisher’, and Roy M. Koerner’ 


The Tambora volcanic signal (acid layer) has been identified in ice cores taken from Agassiz and 
Devon ice caps in the Canadian High Arctic. Oxygen isotope values (representative of annual 
precipitation temperature) and melt percent values (representative of summer temperatures) from 
core sections surrounding the volcanic signal have been examined in detail and compared with 
present day conditions. The results suggest that the Tambora volcanic eruption did not produce 
significant cooling in the Canadian High Arctic. 

On Agassiz Ice Cap the ice representing the year after the volcanic signal shows an increase 
(warming) of both oxygen isotope and melt percent values followed by a return to pre- volcanic 
conditions. On Devon Ice Cap the oxygen isotope values began to decrease (cool) prior to the 
Tambora signal and cool to a minimum 25 years later. Melt percent values on Devon Ice Cap had 
already reached a minimum by the time of the eruption and this persisted for 45 years. 

Based on modern synoptic studies, the circulation pattern during the summer season containing 
the Tambora signal (1816) is best represented by the 1972 analogue. In this analogue a long, 
narrow vortex at 500mb (SO0kPa) extends from the Siberian side of the central Arctic Ocean 
across the Pole deep into Labrador-Ungava, and is held tight against Greenland by a strong ridge 
of high pressure in the Alaska-Beaufort Sea area. This pattern results in strong cold northwesterly 
flow, with frequent light precipitation and very little melt on the ice caps. The pattern is broken 
occasionally by the joining of the Alaska and Greenland ridge which brings clear skies and some 
melt to the islands along the northwestern edge of the archipelago. 


The records from deep ice cores extracted from ice caps in the Canadian Arctic Islands provide 
insight into the climatic conditions in the islands during the decade surrounding the eruption of 
Mount Tambora in Indonesia. It is also possible from these data to address the question of 
whether single volcanic events (such as the Tambora eruption) produce significant deviations in 
proxy annual temperatures and/or proxy summer temperatures in this area of the High Arctic. 
Using modern synoptic-climate analogues, inferences can be made about synoptic circulation 
conditions at the time of the eruption of Mount Tambora. 

Terrain Sciences Division, Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A OE4, Canada. 

Department of Glaciology, Geophysical Institute, University of Copenhagen, Haroldsgade 6, DK-2200, 
Copenhagen N, Denmark. 

Geological Survey of Canada Contribution No. 156789. 


For this study the most complete data are available from a core drilled in 1984 at the top of a 
local dome on the Agassiz Ice Cap (A84) on northern Ellesmere Island (Figure 1). Results from 
another Agassiz Ice Cap core, drilled in 1977, 1.2 km down the flow line from the dome (A77) 
and from a combined time series of three cores taken from the top of Devon Ice Cap (Figure 1) 
in 1971, 72 and 73 (referred to as D123) are also examined. 

__ circle \ 


Figure 1: Location of ice cap deep core drill sites in the Canadian Arctic Islands and Greenland. 

These cores have been extensively discussed elsewhere (Paterson et al. 1977; Koerner 1977; 
Koerner and Fisher 1981; Fisher et al. 1983; Koerner and Fisher 1985; Fisher et al. 1985; Alt 
1985; Alt et al. 1985; Fisher and Koerner 1988). Here we will confine ourselves to the 
parameters and analyses which come closest to providing climatic data with an annual resolution. 
Every attempt has been made to provide an accurate annual time scale and to make these 
consistent for the period surrounding the Tambora eruption. It should be noted that northern 
hemisphere eruptions deposit acidic aerosols on the snow within months of the event, but 
southern hemisphere acid signatures first appear as much as one year after the southern eruptions. 


Melt Features and Oxygen Isotopes 

Melt layers (ice formed by the refreezing of meltwater) can be identified in ice cores by their 
relatively low concentration of air bubbles. In the upper reaches of an ice cap this melt is 
indicative of summer warmth (Koerner 1977). It is expressed as either melt-layer thickness (m) 
or as a percent of the total annual accumulation (PC). As melt (m or PC) can never be less than 
0, possibly very severe summers are not adequately represented in the melt record. Table 1 gives 
the size of errors associated with time series of PC. 

5(8O) is the '*O/"*O ratio expressed as the fractional difference between the ratio in the sample 
and the ratio in "standard mean ocean water" (SMOW) measured in percent. In polar snow, 6 
is negative. Initially, 6 was used as an indicator of mean annual temperature due to its dependence 
on the temperature at which condensation takes place. However up to seven non-temperature 
effects can alter the 6 at a given site (Dansgaard et al. 1973; Fisher 1979; Fisher and Alt 1985; 
Johnsen et al. 1989). For the present discussions the 6 values should be viewed as representing 
the mean annual precipitation temperature (temperature during precipitation events). Table 1 gives 
the size of the errors associated with time series of 6. 

The Tambora Volcanic Signal 

The volcanic peaks are identified by measuring the electrical conductivity (ECM) of core 
segments using brass electrodes with a 1250 DCV potential between them. The resulting values 
are plotted on a time scale derived from models and measurements of annual layer thickness 
(Koerner and Fisher 1985). Major acid layers are correlated with those in the absolutely dated 
Dye 3 core (Hammer 1980). The Tambora volcanic signal appears in the Dye 3 core in 1816. 

In the A84 core the Tambora signal was identified in core 16 (Figure 2, bottom left). This core 
segment is in the firn at a depth of 27 m, is 146 cm long and represents approximately 16 years 
of accumulation. The core segment containing the peak signal from the eruption of Mount Laki 
in Iceland which occurs at 30 m depth in the ice is shown in Figure 2, bottom right. The peak 
ECM value for this Icelandic volcano is much higher than that from the Indonesian volcano, 

The melt-layer thickness values m for these cores have been plotted in a manner consistent with 
the ECM values. The melt and acid feature data are lined up within 5 cm. The Tambora event 
falls between two melt features whereas the Laki acid layer comes 40 cm above the big melt 
feature in core 18 (Figure 2, top). 

Time Series of Acid, Melt and Oxygen Isotope Values for Agassiz 

The annual average values of ECM, PC and 6 have been plotted on the volcanic time scale for 
the Agassiz 84 core (Figure 3). This is probably the best time scale of the various Agassiz Ice 
Cap cores, and great care has been taken to align the three stratigraphies, however discrepancies 
of a year are possible. The Laki signal is absolutely dated as 1783 and the Tambora signal as 

Both the 6 and PC values (Figure 3) reach a maximum just following Tambora. The average 
annual 6 has been calculated for the period of 1941-70, which is used as a meteorological normal. 
Compared to this, the 1816 6 value is slightly above the modern normals. The 1941-71 melt 
normals could not be calculated for A84. Instead the mean PC (dotted line, Figure 3) and the 



a oF ho rae) 5 


Figure 2: 

CORE 15 > CORE 16—>| 

Melt-layer thickness (m) shown as actual observations of the thickness of individual layers of 
ice (top) and volcanic electrical conductivity measurements ECM values from the Agassiz 84 
core segments containing the Tambora and Laki volcanic signals (bottom). Values are plotted 
on a depth scale with top of the core to the left. More than one melt layer can occur in a year. 
Core 16 is 146 cm long, so 5 cm is equivalent to the smallest horizontal increment on the ECM 


Table 1: Errors in Percent Melt PC and Oxygen Isotope 6 Time Series. 

Site Interval Start AD PC SD PC i) SD 6 Accumulation 
years average noise average noise (ice) 
lyr 5S yr 1 yr 5 yr cm/yr 
% % %o % ‘h 
A77 500 1946 2.8 Or 2-514) S15 0:48: O:335=— 17:5 
A84 800 1961 4.1 >25 65.5) -28:5 0:32,10:23 9.8 
Devon' 500 1956 7.0 >S, 186) -28.:0 0.55 0.40 23.0 

' Devon combined record; 6(73 +72) and PC(71+72+73). 
Note: SD is the standard deviation. The Devon SD(noise) data has been measured, but the A77 and A84 noise data 
is estimated. 

Eureka mean July temperature (5.7°C) for the 1951-60 decade have been calculated. The 1941-70 
Eureka July normal temperature (5.4°C) is slightly colder than the 1951-60 decade. The 1816 
melt is also below the 1951-60 decade mean or probably near the modern normal, whereas the 
1817 melt is considerably greater - comparable to the 1951-60 warm period. 

When the A84 values are plotted as five-year averages (Figure 4) the 6 profile might well be 
interpreted as showing a cooling immediately following Tambora. This is in direct contrast to the 
annual averages that show an immediate warming. Care must be used, therefore, in interpretation 
of averaged values in studying the short-term effects of single volcanic eruptions. 

As mentioned, the A84 core has the most accurate time scale but it is at the top of a dome. Here 
the light winter snow is consistently scoured (i.e., blown away). This results in a 6 record which 
is "warmer" than it would be if the winter snow was included. The A77 core lies sufficiently 
downslope from the dome to escape the scouring effect. A detailed plot of annual average 6 
values for the Tambora period from A84 and A77 (Figure 5) shows that the minimum preceding 
the Tambora signal is much colder in the unscoured core than in the scoured (A84) core. Based 
on the most recent time scale for the A77 core, this puts 1816 near the bottom of this minimum 
followed by a rise to 1970 normal values by 1819. 

Time Series from Other High Arctic Cores 

From the Devon blended 1971, 72 and 73 core records (D123) only five-year averages of 6 and 
melt percent are available (Figure 6). The time scale is the vertical velocity time scale fine-tuned 
by analysis of annual layering as deduced by seasonal swings in microparticle concentrations and 
radiocarbon dating of gas bubbles in the ice (Paterson et al. 1977) and corrected for the location 
of the well-marked Laki eruption in the meltwater electrolytic conductivity records (Koerner and 
Fisher 1981). It is accurate to within a few years at the 1816 level. The 1941-70 normals are 
shown for both 6 and melt percent (Figure 6). 


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1940 1920 1900 1880 1860 1840 1820 ' 1810 1800 

vt oO 
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Figure 3: Annual averages from the Agassiz 84 core (A84): (a) oxygen isotope values 6; (b) melt 

expressed as percent of accumulation PC; and c) volcanic ECM values. The arrows on the 
ECM plot show the magnitude of the actual measured peak ECM value for various volcanic 
signals. The dotted line shows the 1951-60 PC decade average. The dashed line shows the 

1941-70 6 30-year average. 


8 0 (%) 

Melt (°/) 


-26 - 26 
-27 97 
-28 -28 
-29 5 8) 
40 cs 
SZ o S x 
2 8 % 30 
30 & E 
© nd 
0 0 
oO 1940 1920 % 1900 BS 1860 1840 © 1800 2 1760 
2 2 © 2 rs 

Figure 4: Five-year averages from A84 of oxygen isotope 6 (dashed line 1941-70 average) and melt 
percent PC. 

The most striking feature in the Devon Ice Cap blended record is the very cold summer period 
indicated by consistently low melt values during the whole period 1810-55. The eruption of 
Tambora occurred well after the beginning of this period. The Laki eruption, on the other hand, 
occurred during a period of increasing melt which reached almost to present normal values by 

On the 6 plot 1816 falls on a cooling trend which began before 1810 and reaches its lowest values 
for the 200-year period during the 1830s. 

Comparing the 10-year melt averages for A77, D123 and Dye 2 on Greenland (Figure 7), we see 
that 1816 falls in a period of generally cold summers at all sites. In all cases the cold period 
began before the eruption of Tambora. At Dye 2 the lowest melt values occur in the 1820s and 
30s but the long flat cold period of the D123 cores is not present. 

The Effect of Single Volcanic Events on High Arctic Climate 

None of the ice core records examined above shows definitive evidence of cooling resulting from 
the eruption of Mount Tambora. Those records which reach a minimum at some time following 
1816 all show a cooling trend beginning before the Tambora volcanic signal. The same is true 
of the Laki eruption. Two other volcanic signals have been identified in the A84 core (Figures 
3 and 4), dated by correlation with the Dye 3 volcanic record and also plotted on the D123 


6'°0 (%o) 

a @ 
A77 2 ca 
- =P 252 Fz 

MELT (%) 


© 1810 

Figure 5: Detailed comparison of the annual mean A84 and A77 oxygen isotope values 6 and the A84 
melt percent values PC from around the Tambora volcanic signal. The five-year mean values 
(solid lines) and the 30-year 6 normals from 1941-70 (dashed lines) are shown for the oxygen 
isotope values. The dotted line shows the average PC values for 1951-60. 


record (Figure 6). The year of the eruption of Mount Agung in Indonesia is also shown on these 
figures. The Agung signal has not been positively identified in the Canadian cores as it was not 
sufficiently acidic. The five-year A84 PC averages (Figure 4) suggest cooling following Katmai, 
but close examination of the annual melt record (Figure 3) shows the season prior to Katmai was 
also cold. Both melt and 6 values show high values following the Krakatau signal. On Devon Ice 
Cap the five-year averages for Krakatau drop sharply in summer melt but the 6 values are already 
low. The Agung eruption appears to occur at the bottom of a 6 minimum in the A84 core. Melt 
data are not available past 1961. In the Devon cores the eruption follows a 6 minimum and is on 
a well-established downward melt trend. 

- = 
-27 =Pi7/ 
 -28 f-28 
| lee" 
Ma o 
< 20 E -26 
=< oO 
— < 
: | 
10 10 

1980 & 1940 © 1900 1880 1860 1840 © 1800 2 
2 2 2 = 

Figure 6: Five-year averages for D123, the Devon blended record (1971, 72 and 73 cores), of oxygen 
isotope 6 and melt percent PC. The 30-year normals, 1941-70 are indicated (dashed lines). 

The very cold summer of 1964 in the Canadian High Arctic, and the subsequent generally lower 
summer temperatures have been attributed to the effects of dust from Agung (Bradley and 
England 1978). Close examination of the hemispheric temperature plots of Dronia (1974) and 
Kelly et al. (1982), Figure 8, show that in both cases the hemispheric temperatures had begun 
to cool long before the eruption of Agung in 1963. The rather dramatic drop of July mean 
temperatures seen in the plots from the northern Canadian Arctic Island stations (Figure 9) is, 
in fact, a result of the record high temperatures in the 1962 season. 


These results do not appear to indicate that single volcanic eruptions cause lower summer or 
annual temperatures in the northern Canadian Arctic Islands. This does not rule out the possibility 
that multiple eruptions in a period could have a cumulative effect on temperatures (Hammer et 
al. 1980) or that single volcanic events produce abrupt, short-lived temperature depressions on 
a hemispheric scale (Bradley 1988). Single events could also be responsible for significant 
anomalies in the atmospheric circulation regime in the Canadian Arctic Islands such as occurred 
in the summer of 1964 (Alt 1987). 

Summary of Core Results 

Now we can review what the ice core analyses reveal about climate in the area at the time of the 
Tambora volcanic signal. The results are expressed in Table 2 as simple estimates of the 
temperature anomalies with respect to the modern normals (1941-70). 

On the Agassiz Ice Cap the summer melt conditions, and thus the summer temperatures, were 
near or slightly below the 1941-70 normals. There was a rise in the melt values immediately after 
the Tambora event to values similar to the relatively warm 1951-60 decade as experienced at 

The annual temperature (or more accurately the annual precipitation temperature) on Agassiz Ice 
Cap appears to have been lower than the 1941-70 normals. The scoured A84 core shows the 
Tambora signal to be part of a slight rise from below-normal conditions to above 1941-70 
normals. The unscoured core A77 shows 1816 to be on a warming trend from a very cold period. 

On Devon Ice Cap there was very little summer melt, indicating very cold conditions. These 
conditions began around 1810 and persisted until the late 1850s. This is the longest very cold 
period in the 800-year record. 

On Devon Ice Cap the Tambora signal falls on a cooling trend of annual (or precipitation) 
temperature beginning about 1810, when the oxygen isotope values were very near the modern 
normals. This cooling trend could be viewed as part of a general decline beginning before the 
time of the Laki signal. 

Table 2: Conditions on Canadian Arctic Island Ice Caps During the Period of the Tambora Volcanic 
Signal as Deduced From Ice-Core Records. 

Season Ice Cap Temperature Remarks 
SUMMER Agassiz 0 (normal) then rising 

(from melt 

percent records) Devon - - (very cold) already very cold 
ANNUAL Agassiz - (cold) on a warming trend 
(from oxygen 

isotope values) Devon - (cold) on a cooling trend 


PC 10yr 







8 $8 188 1S@ 288 288 

AD 1971--< 

Figure 7: Comparison of 10-year averages of A77, Devon blended D123 and Dye 2 (Greenland) melt 
percent values. 


Synoptic Conditions 

Based on the core results for the period around the Tambora signal it is now possible to examine 
the synoptic circulation patterns which would be expected to produce these conditions on the two 
ice caps. Previous studies of synoptic analogues and ice-core results (Alt 1985; Alt et al. 1985; 
Alt 1987) have suggested that this period of the Little Ice Age was dominated by summers similar 
to the summer of 1972 (Figure 10). The most important feature of the 1972 circulation analogue 
for the study area is the persistence of a long deep 500mb (50kPa) vortex held against Greenland 
by a strong ridge over Alaska and western Canada. This produces persistent northwesterly flow 
into the Canadian Arctic Islands from the central Arctic Ocean and a deep layer of very cold air 

in the northern Baffin Bay area. : 

a) °C 




1880 1900 1920 1940 1960 1980 

Figure 8: Two depictions of the annual temperature record for the arctic: (a) annual temperature 
departures from the 1946-60 reference period for 65-85°N (after Kelly et al. 1982); and 
(b) annual deviations from the 25-year mean 1949-73 of thickness of the 500/1,000mb layer for 
65-90°N (after Dronia 1974). 




7 =| 
> | 
ro) -| 
St + 

-2 + - + + + + + + + + + + + + 

I950 52 54 56 58 60) 62, 645566 68 70 72 74 76 

Figure 9: Normalized deviation from the mean of July temperature for Canadian Arctic Islands stations 
[July mean - July normal)/July standard deviation] from P. Schofield (personal 

These features are evident from comparison of the mean July 500mb (S0kPa) height contours for 
the period 1948-78 with those of 1972 (Figure 11a,b). We see that the 1972 vortex is deeper and 
narrower than the mean, and shifted eastward from the mean position by a ridge over the 
Beaufort Sea. The flow into the High Arctic Islands is stronger than normal as seen by the closer 
spacing of the contour lines. There is also a strong ridge over the Barents Sea. The mid-latitude 
circulation during the entire 1971-72 season was stronger than usual and distinctly meridional 
(i.e., with strong north-south components). 





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500 mb and surface 

Figure 10: Schematics of summer synoptic characteristics for the Agassiz and Devon core-site area for 
various periods of the last 800 years. These were deduced by applying modern analogues to 
the ice-core results; represented here by the five-year averages of oxygen isotope 6 and melt 
percent PC from Devon Ice Cap (Alt 1985). 


Figure 11: Mean July pattern of 500mb (SOkPa) height contours in decameters (dam). In order to focus 
on the polar vortex, contours up to 560 dam only are shown: (a) for 1948-78 (after Harley 
1980); and (b) for 1972. 


In order to examine the surface weather conditions associated with the 1972 anomaly, the actual 
synoptic charts for 3 July 1972 are shown in Figure 12a,b. Here we see the long upper vortex 
(Figure 12a) with multiple centres, one over Labrador-Ungava, a second in northern Baffin Bay 
and a third over the Pole. This supports a surface trough (Figure 12b) across northern Ellesmere 
Island down the west coast of Greenland to Davis Strait. The strong northwesterly flow, indicated 
by the closely packed isolines, extends across the Canadian Arctic Islands into Keewatin, Hudson 
Bay and James Bay at all levels from the surface to 500mb (about 5,000m). Cold moist air from 
the central Arctic Ocean is pushed south into these areas. In the northern islands extensive low 
cloud, high humidity and frequent light precipitation accompany these conditions. Over Devon 
Ice Cap precipitation may be enhanced by the persistence of the northern Baffin Bay low, which 
picks up additional moisture from the open water. 

Both the mean July 1972 (Figure 11b) and 3 July 1972 (Figure 12a and b) patterns appear to be 
consistent with conditions proposed for the Hudson Bay region during 1816. Wilson’s (1983) 
studies show prevailing NW-NE winds in June and July 1816 and also suggest as a modern 
analogue the summer of 1972. High pressure west of Hudson Bay (in the case shown here, an 
extension of the Alaska ridge) is an important feature of her proposed 1816 circulation patterns. 

Figure 12 (a): Synoptic chart for 3 July 1972: 500 mb (50 kPa) height contours in decametres (dam). 


/ 1830 C 

Figure 12 (b): Synoptic chart for 3 July 1972: surface pressures in mb. 

Figure 12a,b also gives us an indication of the synoptic conditions that previous detailed synoptic 
studies of summers from 1960-78 have shown could produce melt on Agassiz Ice Cap and not 
on Devon Ice Cap. Close inspection shows that the ridge extending from the Beaufort Sea to 
southern Manitoba (which is responsible for keeping the trough or vortex tight against Greenland) 
can be seen pushing northeast across the northern Canadian Arctic Islands toward the north 
Greenland ridge. If these join (as they did later in July 1972), the vortex is cut off, becoming a 
closed low. The closed system can lie over Labrador-Ungava or farther west, as happened in July 
1972. Along the northwestern edge of the islands the ridge results in subsidence through the 
whole troposphere, which dissipates the cloud and fog. Melt is produced by increased solar 
radiation (sometimes aided by warm-air advection) over the northern and western ice caps. Devon 
Ice Cap, however, is often under the influence of the cyclonic circulation in Baffin Bay and not 
as likely to experience melt. In fact, on Devon Island summer accumulation (snow) may occur 
under these conditions. These ridging conditions are often brief but they can produce significant 
melt on the ice caps and a touch of summer in the northwestern islands, as happened in mid-July 
1972 (Alt 1987). 

We can also say that this pattern resembles the mean winter conditions, and suggest that in years 
of this kind the winter circulation is never really broken down. The temperature gradients remain 
strong as do the mid-latitude westerlies. Summer comes only briefly to the ice caps if and when 
the blocking ridges join across the northern islands. These ridging conditions are more effective 
in the northwestern islands and may not produce any melt at the core site on Devon Ice Cap. 



The Tambora signal can be identified as an acid layer in the Agassiz 84 core. The oxygen isotope 
5 and melt values PC, even allowing for a one-year discrepancy and other considerations such 
as noise and scouring, do not show evidence of cooling due to the eruption of Mount Tambora, 
although it may have occurred part way down a cooling trend. Nor is there definitive evidence 
of cooling in the northern Canadian Arctic Islands following the eruption of Laki, Krakatau, 
Katmai or Agung. 

On Agassiz Ice Cap, conditions in the year dated as 1816 were near, or somewhat below, modern 

normals (1941-70) but rise to a secondary peak immediately following the Tambora signal. Care - 
must be taken when interpreting average values for periods longer than a year as they can easily 

obscure the short-term variations. However, on the Devon Ice Cap blended five- year average 

plots, 1816 falls in a prolonged period of very low summer melt and below modern normal 6 

levels (annual or precipitation temperature); both of which began about 1800. Similarly the Dye 2 

10-year average plot shows Tambora occurring on a well established cooling trend. 

The climatic conditions suggested by the ice-core analyses around the Tambora eruption strongly 
resemble those of the summer of 1972, which has been identified as the modern analogue for 
melt suppression on High Arctic ice caps. This pattern, which features a long deep upper vortex 
extending from the Siberian side of the central Arctic Ocean across the eastern Canadian Arctic 
Islands to Labrador-Ungava and a strong ridge from the Beaufort Sea and Alaska into the 
prairies, appears to be compatible with synoptic interpretations from other parts of Canada. This 
pattern represents an intensification of the conditions which appear to have dominated the latter 
part of the Little Ice Age in the High Arctic islands. 


Alt, B.T. 1985. A period of summer accumulation in the Queen Elizabeth Islands. Jn: Critical 
Periods in the Quaternary Climatic History of Northern North America. Climatic Change 
in Canada 5. C.R. Harington (ed.). Syllogeus 55:461-479. 

. 1987. Developing synoptic analogues for extreme mass balance conditions on Queen 
Elizabeth Island ice caps. Journal of Climate and Applied Meteorology 26(12):1605-1623. 

Alt, B.T., R.M. Koerner, D.A. Fisher and J.C. Bourgeois. 1985. Arctic climate during the 
Franklin Era as deduced from ice cores. In: The Franklin Era in Canadian Arctic History. 
Pat Sutherland (ed.). National Museum of Man, Mercury Series, Archaeological Survey of 
Canada Paper 131:69-92. 

Bradley, R.S. 1988. The explosive volcanic eruption signal in northern hemisphere continental 
temperature records. Climatic Change 12:221-243. 

Bradley, R.S. and J. England. 1978. Volcanic dust influence on glacier mass balance at high 
latitudes. Nature 271:736-738. 

Dansgaard, W., S.J. Johnsen, H.B. Clausen and N. Gundestrup. 1973. Stable isotope glaciology. 
Meddelelser om Gronland 197(2):1-53. 


Dronia, H. 1974. Uber Temperaturanderungen die frier Atmosphare auf der Nordhalbkugel in 
den letzten 25 Jharen. Meteorologische Rundschau 27:166-174. 

Fisher, D.A. 1979. Comparison of 10° years of oxygen isotope and insoluble impurity profiles 
from Devon Island and Camp Century ice cores. Quaternary Research 11(3):299-305. 

Fisher, D.A., R.M. Koerner, W.S.B. Paterson, W. Dansgaard, N. Gundestrup and N. Reeh. 
1983. Effect of wind scouring on climatic records from ice-core oxygen-isotope profiles. 
Nature 301(5897):205-209. 

Fisher, D.A., R.M. Koerner, N. Reeh and H.B. Clausen. 1985. Stratigraphic noise in time series 
derived from ice cores. Annals of Glaciology 7:76-83. 

Fisher, D.A. and B.T. Alt. 1985. A global oxygen isotope model - semi empirical, zonally 
averaged. Annals of Glaciology 7:117-124. 

Fisher, D.A. and R.M. Koerner. 1988. The effects of wind on 6(O"*) and accumulation give an 
inferred record of seasonal 6 amplitude from the Agassiz Ice Cap, Ellesmere Island, 
Canada. Annals of Glaciology 10:34-38. 

Johnsen, S.J., W. Dansgaard and J.W.C. White. 1989. The origin of Arctic precipitation under 
present and glacial conditions. Tellus Series B, 41B(4):452-468. 

Hammer, C.U. 1980. Acidity of polar ice cores in relation to absolute dating, past volcanism, 
radio-echoes. Journal of Glaciology 25(93):359-372. 

Hammer, C.U., H.B. Clausen and W. Dansgaard. 1980. Greenland Ice Sheet evidence of 
post-glacial volcanism and its climatic impact. Nature 288(5788):230-235. 

Harley, W.S. 1980. Northern hemisphere monthly mean 5O kPa and 100 kPa height charts. 
Environment Canada, Atmospheric Environment Service CLI-80:29. 

Kelly, P.M., P.D. Jones, C.B. Sear, B.G. Cherry and R.K. Tavakol. 1982. Variations in surface 
air temperatures. Part II: Arctic regions, 1881-1980. Monthly Weather Review 

Koerner, R.M. 1977. Devon Island Ice Cap: core stratigraphy and paleoclimate. Science 

Koerner, R.M. and D.A. Fisher. 1981. Studying climatic change from Canadian High Arctic ice 
cores. In: Climatic Change in Canada 2. C.R. Harington (ed.). Syllogeus 33:195-215. 

. 1985. The Devon Island ice core and the glacial record. In: Quaternary environments; 
eastern Canadian Arctic, Baffin Bay and western Greenland. J.T. Andrews (ed.). Allen 
and Uwin, Boston. pp. 309-327. 

Paterson, W.S.B. and seven others. 1977. An oxygen-isotope climate record from Devon Island 
Ice Cap, Arctic Canada. Nature 266(5602):508-511. 

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a{i> a3! ii 


Europe (including Iceland) 

j - 
zal sp ligel rat eee 



1816 - a Year without a Summer in Iceland? 

A.E.J. Ogilvie’ 


There has been considerable speculation as to whether the eruption of Mount Tambora in April 
1815 caused a world-wide lowering of temperatures and a “year without a summer" in the 
following year of 1816. In this paper, the weather during 1816 is detailed for one specific 
location: Iceland. The weather data used are taken from documentary accounts written at 10 
different sites in Iceland. These suggest that the winter and spring of 1816 were very cold and 
unfavourable in most parts. The summer was mainly cold in the north, wet in the east and highly 
variable elsewhere. Many accounts of the autumn focus on the variability of the weather. 
Although it would seem that, on the whole, the summer weather was not sufficiently extreme for 
this year to be termed a “year without a summer," adverse weather did cause some impact on 
society. It seems very likely that there was direct climatic impact on important agricultural 
practices such as the hay harvest and the growing of vegetables. The Arctic sea ice, although not 
unusually heavy or prolonged in 1816, had a direct impact in northern Iceland, hindering fishing 
and sealing. Indirect impacts on society are less easy to establish. However, it seems likely that 

some social stress described in 1816 may be at least partly attributed to the climate. 


Although the precise nature of the effects of volcanic eruptions on the general circulation of the 
atmosphere are, as yet, unknown, there can be little doubt that major volcanic eruptions do affect 
the Earth’s climate (e.g., Lamb 1970; Kelly and Sear 1984; Sear et al. 1987; Bradley 1988). The 
possible effects of one very large eruption - that of Mount Tambora in April 1815 - has excited 
particular interest. Although some researchers (e.g., Landsberg and Albert 1974) have concluded 
that this eruption did not have significant climatic effects, others have provided convincing 
evidence to show that the subsequent year, 1816, was anomalously cold in many places (Stothers 
1984; Kelly et al. 1984). The year 1816 has even been termed the "year without a summer" 
(Stommel and Stommel 1979). 

In this paper, the weather during 1816 is considered for one specific location - Iceland. In order 
to place the year in context, the general climate of Iceland is considered first, both for the 
twentieth century, and in terms of climatic variations in the past. Possible climatic impact in 
Iceland during 1816 will also be discussed. ; 

The Present and Past Climate of Iceland 
The Twentieth Century Context 
Our knowledge of the climate of Iceland is derived from two main data sources. The principal 

of these is modern instrumental data. By the late nineteenth century, around 20 observing stations 
were in existence, and with the establishment of the Icelandic Meteorological Office (Vedurstofa 

' Climatic Research Unit, School of Environmental Sciences, University of East Anglia, Norwich NR4 7TJ, U.K. 


Islands’) in 1920, the number of stations grew. By 1955, there were 66. From 1966 onwards, 
the number has varied between 120 to 130 (Einarsson 1976, pp. 12-13). Information from these, 
and other observing stations in the North Atlantic and Polar regions, plus oceanographic data, 
has enabled a general picture of key factors in the climate and weather of Iceland to be 
established. These are summarized below. For more detailed discussions on this topic, see 
Eythorsson and Sigtryggson (1971) and Einarsson (1976). 

Main Features of the Climate of Iceland 

The principal features of Iceland’s climate are determined by its location at the frontier zone of 
two very different air masses; cold polar air from the north, and warmer maritime air from the 
Atlantic. Depressions moving toward Iceland from the western Atlantic often slow down as they © 
near the southwestern corner of Iceland, thus maintaining a flow of mild Atlantic air over the 
country. This process causes thaws in winter, and rain and cool temperatures in summer. When 
these depressions cross Iceland and move toward Norway, a flow of polar air may take their 
place and bring much colder weather, especially in the northern part. 

The alternating cold and milder air masses that Iceland experiences at varying intervals, and for 
different durations, are the prime cause of the variability of Iceland’s climate. This variability is 
exacerbated by the two major ocean currents which flow around the island; the cold East 
Greenland polar current, and the warmer Irminger current. The Arctic drift ice also has 
considerable influence on the climate of Iceland. Most noticeably, when the ice is present off the 
coasts, both land and sea temperatures are lowered. 

Although weather conditions in Iceland vary greatly, generally winters are mild compared with 
other northern continental locations, and summers tend to be cool. Typical temperature ranges 
during the winter months of December, January and February vary between -2 and 1°C. The 
warmest summer month is generally July, with a mean temperature varying from around 8 to 
11°C, depending upon location. 

The Past Climate of Iceland: Introduction 

Information about the climate of Iceland derived from modern data is augmented and amplified 
by what is known of the past climate of Iceland. This is derived largely from documentary, 
historical evidence: the nature and use of such evidence is discussed briefly below. Although we 
cannot hope to gain as accurate a picture from documentary evidence of climate as from modern 
instrumental data, such evidence can act as a guide to what may have occurred in the past when 
no other data are available. To this end, proxy temperature variations based on the use of 
historical documentary evidence have been derived by Bergthdérsson (1969) and by Ogilvie 
(1984a, 1986, 1990). Incidence of the sea ice off the coast of Iceland in the past has been 
estimated by these same authors, and by Koch (1945). 

In the sections below, probable variations in the past climate of Iceland from medieval times to 

the early nineteenth century are outlined. Prior to this, the available data sources for this period 
are discussed. 

! The Icelandic characters "pb" and "4" (for "th") and all accents are retained wherever these are used in the original. 


Data Sources 

The accuracy of any proxy-temperature indicator will depend on the quality of the data used. To 
ensure high quality of documentary evidence, all sources must be analyzed carefully in order to 
establish their reliability. Key questions to ask here are whether the author was close in time and 
space to the events described; if this is the case, then a source is much more likely to be reliable 
than if he were not. For more detailed discussions on source analysis in general, see Bell and 
Ogilvie (1978) and Ingram et al. (1978). For discussions of the analysis of Icelandic sources see 
Vilmundarson (1972) and Ogilvie (1981, 1984a, 1990, 1991). 

Iceland’s climatic history may be traced back to early settlement times in Iceland (from about 
A.D. 870 onwards). However, the quality and availability of climatic and weather data vary 
considerably. For the period up to about 1170, there are no contemporary documentary sources, 
and only brief and sporadic comments on weather and climate may be found in existing sources. 
For the thirteenth century, a few reliable sources give some indication of possible changes in 
climate. Many more descriptions of weather and climate exist for the fourteenth century. The 
fifteenth century and the first half of the sixteenth century are very poorly documented. Typical 
sources for this period are certain sagas, the medieval annals and works of geographical 

From the early seventeenth century onward, many more reliable documents become available. 
For the early to mid-eighteenth century, there is extensive coverage from a variety of different 
sources including annals, travel accounts, government reports and weather diaries. These give 
information for most seasons in many different parts of Iceland. For the late eighteenth and early 
nineteenth centuries, sources of climatic and weather information are very full and detailed. 

The earliest quantitative observations taken in Iceland date from the mid-eighteenth century 
(Eyb6rsson 1956; Kington 1972). However, these, and subsequent late-eighteenth and nineteenth- 
century observations only cover a few months or years. Continuous temperature observations 
commence in 1846 (Sigftisd6ttir 1969). These were made at Stykkishélmur, in the west. For the 
period 1820-54, observations of temperature were taken in Reykjavfk or the near vicinity by Jén 
Porsteinsson (1794-1855). A part of this important series was subsequently lost for many years. 
However, the missing data were recently found by Trausti Jénsson of the Icelandic 
Meteorological Office, and he is engaged in their analysis (J6nsson, personal communication). 

The Climate of Iceland from Settlement Times to about 1600 

Iceland was settled, primarily from Norway, in the late ninth and early tenth centuries. 
Circumstantial evidence suggests a fairly mild climate around this time. A cold period may have 
occurred from about 1180 to 1210, while from about 1211-32 the climate may have become 
milder. An early geographical treatise written in approximately 1250 (The King’s Mirror) 
mentions much sea ice between Iceland and Greenland at this time, and refers to Iceland’s cold 
climate. However, it is difficult to draw firm conclusions from statements such as this. From 
about 1280 to 1300 the climate seems to have been fairly cold. During the early years of the 
fourteenth century, severe weather is mentioned only infrequently (in 1313, 1320, 1321 and 
1323), so this period may have been mild. Milder weather may well have continued to past the 
mid-fourteenth century. The years 1360-80 are likely to have been colder. Little information is 
recorded from the 1380s. Only two severe years are noted for the 1390s. Evidently, 1412-70 was 
mild, and the 1480s or 1490s were years of dearth, possibly caused by severe weather. However, 


very little information is available for 1430-1560. Likely the latter part of the sixteenth century 
was mainly severe. A detailed discussion of all medieval historical sources containing comments 
on the weather and climate, together with an analysis of their evidence, may be found in Ogilvie 

The Climate of Iceland from 1601 to about 1850 

The first and second decades of the seventeenth century were, overall, probably relatively mild. 
The years 1620-40 were cold, but 1641-70 was distinctly mild. From 1671-90, temperatures were 
colder. The 1690s were very cold. The early years of the eighteenth century were relatively mild, 
especially the first decade. The 1730s, 1740s and 1750s were cold, especially the two latter 
decades. The 1760s were somewhat milder, the 1770s cooler again. The period 1781 to 1820 was | 
cold on the whole. The year 1816 must be assessed in the context of this prevailing background, 
with mainly cold conditions spanning most of the preceding four decades. From 1821 to 1841 
the climate is likely to have been milder, while the 1840s were very mild. Fora fuller account 
of climatic variations in Iceland during the seventeenth and eighteenth centuries, see Ogilvie 
(1981, 1984a, 1986, 1990). 

The Weather in Iceland During 1816 

Data Sources 

In order to build up a clear picture of weather and climate during 1816, a number of sources 
were Selected for detailed analysis. The main sources used are letters written by "Sheriffs" or 
government Officials in the 20 or so different districts or Sys/a (plural Sysslur) of Iceland. These 
letters contain information on such topics as grass growth and hay crop, trade, health and disease. 
They also report on the weather, sometimes in very great detail. One letter used here (from 
Snefellsnessysla, in the west) gives daily data, as well as seasonal summaries. The letters were 
sent at least annually, sometimes more frequently, to the Danish government in Copenhagen. 
Written in Danish, they are all unpublished, and are now kept in the National Archives in 

For this discussion of 1816 weather, letters were chosen from nine different sites in Iceland. Use 
was also made of one other source; an annal, called Brandsstadaandll. This was written by Bjorn 
Bjarnason (1789-1859) at Brandsstadir, Bl6ndudalur in Austur-Hunavatnssysla in the north. This 
annal describes events that occurred in Iceland each year from 1783 to 1858, and includes 
detailed weather descriptions. Other available sources were not used. It was felt, however, that 
the sources used here were adequate to provide good regional coverage over the year. As 
Iceland’s climate is regionally quite variable (Eythérsson and Sigtryggsson 1971; Ogilvie 1984a) 
it was essential to consider different parts of Iceland. 

The sites at which these various sources were written are shown in Figure 1. The sources are 
from (in the order given on the map): (1) Ketilsstadir in the district of Sudur-Mulasysla in the 
east; (2) Gardur in Sudur-Pingeyjarsysla in the north; (3) Médruvellir in Eyjafjardarsysla in the 
north; (4) Vidvik in Skagafjardarsysla in the north; (5S) Brandsstadir in Austur-Hunavatnssysla 
in the north; (6) Gréf in Snzfellsnessysla in the west; (7) Sidumuli in Myrasysla in the west; 
(8) Leird in Borgarfjardarsysla in the west; (9) Reykjavik in Gullbringusysla in the southwest; 
(10) Vfk in Vestur-Skaftafellssysla in the southeast. Some use was also made of a letter written 
at Grund, a site very close to Médruvellir. ' 

' All translations of sources are by the author. 


1816 - The Evidence 


In Table 1, a brief synopsis is given of comments on weather from the sources described above. 
In the left-hand column, the place at which the letter was written is shown according to its 
number on the map. The columns in the centre show the main characteristics of the seasons. The 
term "winter" here refers to the period from mid-October of one year (1815) to mid-April of the 
next (1816). "Spring" covers mid-April to mid-June, "summer" is mid-June to August, and 
"autumn" is September to mid-October. The column on the right-hand side of Table 1 shows 
descriptions of sea ice. 


From Table 1, it may be seen that letters from most districts report a severe winter. Two of the 
letters written in the north, at M6druvellir in Eyjafjardarsysla, and at Gardur in Adaldalur in 
Audur-pingeyjarsysla, stated that the winter was, respectively "more than unusually severe" and 
"very severe." Two other sources, one from V{fk in Vestur-Skaftafellssysla in the southeast, and 
the other from Leird in Borgarfjardarsysla in the west, both noted that the winter was "very 
severe." The writer of the account from Reykjavik wrote that the winter was "severe with much 
snow and frost." 

Some sources stress the variability of the weather this winter. Thus, the letter from Vidvfk in 
Skagafjardarsysla in the north reported that there was "much snow and alternating thaws and 
sharp frosts," and the letter from Sfdumiuli in Myrasysla in the west gives a similar account. 
According to Brandsstadaanndll, the winter was also very severe, but there were some spells of 
calm and good weather in between, for example, from 25 November to 15 December (1815) and 
from 15 January to about 21 February. Interestingly, the letter from Ketilsstadir in Sudur- 
Mulasysla in the eastern part of Iceland stated that the winter was merely average, although there 
was much snow. 


The spring of 1816 was also relatively cold in most districts. The letters written at Vidvfk and 
MOodruvellir in the north characterized the spring as “dry and cold" and "quite severe", 
respectively. According to the latter, the severity took the form of persistent northerly winds, 
frost and cold air. These the writer attributed to the presence of sea ice which lay off the northern 
coasts all spring. The letter from Gardur does not contain a description of spring weather as 
such, but does mention sea ice. This is stated to have been present from the beginning of March 
to mid-June. The spring was said to be “unusually cold" in the letter from Reykjavik. 

According to Brandsstadaanndll, April was severe, but the weather improved at the end of the 
month. The account from Leir4, in the west, also notes severe cold in April, but says that from 
the end of the month to mid-May the weather was mild. From then it became cold again with 
northeasterly winds, sleet and night frost to about 24 June. The Sheriff of Skaftafellssysla, writing 
at Vik, recorded a severe April and a mild May. The first 10 days of June were dry and frosty. 


The weather during the summer of 1816 in Iceland was quite variable regionally. The northern 
sources used make it clear that, in most northern areas, the weather was very poor. Most 
accounts from the south and west report a mixture of both favourable and unfavourable weather. 
The eastern source used here states that the summer was wet. We may look at accounts from 
these regions in more detail, starting with the north. A summary of the data may also be found 
in Figure 2. 


(n & ON Re 

Ket Usstaoiim, SuSur—Mulasysla 
. Gardur, Sudur—Pingey jarsysla 

. Vidvik, Skagafjardarsysla 

6. Grof, Snefellsnessysla 

7. Sidumili, Myrasysla 
Modruvellir, Eyjafjaréarsysla 8. Leird, Borgarfjardarsysla 



. Reykjavik, Gullbringusysla 

Brandsstadir, Austur-Htinavatnsyssla 10. Vik, Vestur-Skaftafellssysla 

Figure 1: Sites of sources used for 1816 weather reconstruction. 


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Vidvik: Dry and cold. Then 
rain set in during harvest 

Grund: Very cold 
(c.mid August) 

Brandsstadir: Dry and 
good during harvest 

Gar6ur: Unpleasant. 

snow, rain and 

Grof: June quite windy 
and rainy. Mainly good 
and calm from end June 
to 14 August. Rainstorm 
on 15 August. Calm again 
through 18 August. Rain 
from 19 August. 


Sidumili: Averagely good eae Very damp and 
escumees rainy 

and n Gh Bae to 24 June. 
Dry during July and August. 
Then became more damp with 
Sly winds to mid October. 

Vik: Heavy rain and storms 

Reykjavik: On the whole, very 11 June to 11 July. 17, 18 
cold andinconstant except for and 19 July Skeidara 

3-17 August when weather mild flooded. 12 July to 11 August 
and good 

dry and good. 12 August to 
12 September rain 

Figure 2: Summer weather reported in Iceland in 1816. 

The Sheriff writing at Vidvfk in Skagafjardarsysla stated that the summer was dry and cold 
except during the harvest around mid-August when there was rain. Stefan P6rarinsson at 
MOdruvellir also mentions long-lasting rain during the harvest. The report from Grund, near 
MOodruvellir, was of a very cold summer. In the letter from Gardur, the summer is said to have 
been unpleasant with alternating snow, rain and frost. The other northern source used here, 
Brandsstadaanndll, disagrees with these accounts. According to this source, the weather became 
good and calm after 25 April, and the summer as a whole was favourable. It should be noted that 
such variability between different areas in Iceland, even sites in close proximity, is not unusual. 
Furthermore, Brandsstadir in Bl6ndudalur, where Brandsstadaanndll was written, is in a fairly 
sheltered location. 

In the east, Sheriff Pall Pordarson Melsted, writing at Ketilsstadir, commented on the damp and 
rainy summer. At V{fk, in the southeast, the summer weather was quite variable. However, this 
variability took the form of quite long spells of fairly stable weather patterns, rather than short- 
term variation on a scale of days. From 11 June to 12 July there were storms and heavy rains. 
Then, from 12 July to 11 August, there occurred "the driest and best weather of the whole 
summer." On 12 August, when the hay harvest had just begun, a rainy period set in and lasted, 
with the exception of a very few days, to September. The rainy weather in the southeast in the 
early part of the summer may have been partially caused by the volcanic eruption that occurred 
in Skaftafellsj6kull. This eruption, and the flood in the river Skeidard, are discussed below. 


Some western accounts of the summer weather show a pattern of variability not dissimilar to that 
noted above for Vfk, although the timing and duration of cold, wet and dry spells is different in 
different locations. At Leir4 in Borgarfjardarsysla, the Sheriff noted that the weather was cold, 
with northeasterly winds, sleet and night frost to about 24 June. Throughout July and August it 
was dry. At Reykjavik, just south of Leird, the report stated that, on the whole, the weather was 
very cold and inconstant except for 3-17 August when the weather was mild and good. The 
account from Grof, in Snefellsnessysla in the west (like Reykjavfk, an exposed coastal site), 
contains daily weather data, and is much more detailed than most of the sources used here. It 
accords quite well with the Reykjavfk report. The daily data given may be summarized as 
follows. On the whole, the month of June was mainly windy with snow, sleet or rain. Ten days 
were characterized as being calm. Only one day of rain is mentioned in July, but many days were 
described as breezy. No storms occurred in July. From 1-14 August, the weather was quite 
favourable. On 15 August and 19-24 August there were rain storms. On 25-27 August there were 
strong south-southwesterly winds with rainshowers, and on 28 August the wind was northeasterly 
with rain and fog. Northeasterly winds continued to the end of the month. One other western 
account, from the inland site of Sfdumuli in Myrasysla, characterized the summer as "averagely 

In spite of the reports from Sfdumuli and Brandsstadir, and some intervals of good weather at 
other sites, when the summer weather of 1816 is considered over Iceland as a whole, it must be 
classed as unfavourable. However, it was not extremely so, and the phrase "a year without a 
summer" does not, therefore, seem appropriate for this year in Iceland. 


Most sources characterize the autumn as mainly stormy and changeable. Thus, in the letter from 
Reykjavik, it is said to have been "stormy all autumn." The report from Gr6f is of "stormy and 
inconstant rainy weather." At Vidvfk, the weather is said to have been wet and inconstant, often 
with strong winds. At Leird, southerly winds are noted. The weather was damp up to mid- 
October. At Vik, September is said to have begun with severe night frost. Subsequently, more 
rain than frost occurred. From 3 to 9 October, there was dry good weather with rime frost. From 
10 October, there were mainly westerly winds with hail, snow, rain, sleet, frost and layers of ice 
on the ground. The autumn weather is said to have been, in general, "unusually changeable." 
According to Brandsstadaandll, there was snow and frost in late September. The first half of 
October was calm and dry, then snows fell. 

Other Environmental Events 

In 1816, a volcanic eruption also occurred in Iceland. This is known from the letter written by 
Sheriff Lydur Gudmundsson at Vik in Vestur-Skaftafellssysla. According to him, the eruption 
began under Skaftafellsj6kull (glacier) some time in May. In June, the eruption was visible over 
16 miles (24 km) away, with an enormous column of rising vapour. "This later divided itself into 
clouds, and caused a bitingly sharp, cold drought until the clouds finally dispersed, and fell as 
a malignant, cold, severe and lasting heavy rain." The eruption does not appear to have had any 
serious effects on the populace, although the vegetable and hay crops were said to have been 
adversely affected. 

Lydur Gudmundsson also reported flooding of the River Skeidar4 on 17, 18 and 19 July. He 
described the river as "flowing out of the bowels of the Skaftafell glacier." Today, the river flows 
adjacent to the neighbouring Skeidardrjékull. This discrepancy may be explained by the fact that 
the glaciers are undoubtedly smaller and of a different shape now than they were in the early 


nineteenth century, and the river is also likely to have changed its course. The Sheriff noted that 
the river flooded a large part of Skeidardsandur (a stretch of sandy plain, washed out from the 
glaciers) and cut off all passage over a much greater distance. Probably the flood was largely 
caused by ice melting during the volcanic eruption. 

Climatic Impact in Iceland in 1816 

The Study of Climatic Impact: Methods and Approaches 

In order to provide an analysis of past events that is as accurate as possible, recent research in 
the field of climatic impact has emphasized the need to adopt a rigorous methodology (Wigley 
et al. 1981; Kates et al. 1985). This is because of the difficulty in isolating and quantifying the | 
effect that climate might have had on society, given all the other social, political and economic 
factors present. Such an exercise is difficult to carry out with present-day data. In the past, when 
fewer economic and climatic data were available, it becomes even more problematic. Although 
the difficulties arising from this can never be entirely eradicated, a number of measures may be 
adopted in order to provide a valid picture of possible climatic impact in the past. Important 
issues to consider before undertaking such a study are: (1) the location of the area to be studied; 
(2) the quality of the available data; (3) the time scale involved; (4) the economy and social 
structure of a given area (e.g., whether primitive or sophisticated); and (5) the strategy, or 
methodology to be employed. These points will be considered further below, with regard to 

Iceland’s Location 

Concerning the first of these points, location, Iceland occupies a marginal area on the borderline 
between environments that lend themselves easily to human habitation, and those that do not 
(e.g., Bergthérsson 1985). The cool climate of the area will obviously be a major factor in 
determining the growth of vegetation of all kinds. The many mountain and cold-desert regions 
mean that any attempts at agriculture will be limited, not only by the climate, but also by the 
amount of land available for such activities. Iceland’s geographical situation thus makes it highly 
suitable for climatic-impact studies. 

Data and Time Scales 

Climatic data available for Iceland during the pre-instrumental era, and the specific sources used 
here, have been discussed above, and their quality established. Regarding the question of time 
scale, the data available make it possible to study climatic impacts in the long term (centuries), 
medium term (years to decades), and short term (a year or less) as it is here. The study of 
climatic impact in the short term has been criticized as giving undue attention to certain crisis 
years (Ingram et al. 1981). However, in this case, it is done within the context of previous 
studies of longer periods (Ogilvie 1981; 1984b). 

Iceland’s Economy and Society 

Before about the mid-nineteenth century, no settlement large enough to be considered a town, or 
even a village, existed in Iceland. There were only isolated farmsteads and a few fishing stations 
on the coasts. The farms were scattered in order to make best use of the land available. 
Settlement was concentrated primarily in the coast and lowland areas of the southwestern, western 
and northern regions. The less hospitable areas of the northeast, southeast and northwest were 
even more sparsely populated. 


Iceland’s economy was based on animal husbandry: the main animals kept were sheep and cattle. 
These provided food in the form of meat and milk products, and also other useful items such as 
wool for clothing. Horses were used for transportation. Fishing was also important, but it was 
not until the twentieth century that this became a major industry. 

During the short summer season, the major task for most Icelanders would be to bring in the 
annual hay harvest - still of great importance today. Hay was grown on the "homefields" (tun), 
near the farm, and on outlying pastures (engi). The hay was given to livestock during winter so 
that they could survive if there was little or no vegetation available. When the weather was 
favourable, certain of the livestock, particularly horses, sheep and gelded cattle, were expected 
to graze outside. These were known collectively as utigangspeningur or “outside livestock." 

From 1380 to the Second World War, Iceland was ruled by Denmark. The trading monopoly 
enforced by Denmark for much of this time frequently worked in Iceland’s disfavour as the 
Danish merchants controlled both prices and the goods available to the Icelanders. 


Throughout Iceland’s recorded history, there are many "crisis-years." These are when the sources 
recount failure of the hay crop, livestock deaths, serious difficulties among the population such 
as the desertion of farms, begging, and even human mortality. Such events invariably occurred 
during very cold years or decades. Because people were so dependent on a successful hay harvest 
for supplementary winter fodder, it appears very likely that a poor harvest or a severe winter 
might have a considerable impact on the populace. 

Rather than take this coincidence of events at face value, however, it is possible to adopt a 
strategy that will help to establish more clearly exactly what was occurring. To this end, this 
possible impact of climate may be divided into direct and indirect impact. 

It is not difficult to demonstrate that climate had a considerable direct impact on biological and 
physical processes; (e.g., on grass growth, hay yield, and on other plants). This may be shown 
Statistically (Ogilvie 1981, 1984b). For example, relationships between temperature, precipitation, 
grass growth and hay harvest may be tested by means of contingency tables (e.g., Table 2). 

Indirect effects of climatic impact include deaths of livestock by starvation (although such effects 
may be compounded by direct impact in the form of cold and damp), and these may also be 
demonstrated by means of contingency tables (Table 3). In both of the tables shown here, the 
results are highly statistically significant. 

Further indirect effects of climatic impact, such as the social problems mentioned above, are far 
harder to prove. Yet frequently much circumstantial evidence is available making it possible to 
show that such effects were very likely to have occurred (Ogilvie 1981). However, in all such 
studies, it is vital to take political, economic and social factors into account, as these invariably 
play a larger role than climate. 

It is not possible to carry out the kinds of statistical tests mentioned above when considering data 
for one year only. However, as climate did have both direct and indirect impacts over longer time 
scales, clearly these would also be felt on an annual time scale. The reality of climatic impact has 
been demonstrated by several researchers using both modern and historical data (e.g., 
Bergthérsson 1966, 1985; Fridriksson 1969, 1972; Bergthdérsson et al. 1988). 


Table 2: Summer Temperature and Grass Growth in Iceland 1601-1780'. 






' chi? = 84.0 

Summer Temperature 















72(31.6 %) 

228(100 %) 

Table 3: Winter Temperature and Livestock Deaths in Northern Iceland 1601-1780. 



Poor condition/disease 

Good condition 


' chi? = 62.8 

Winter Temperature 




















In the section below, climatic impact during 1816 is considered. The direct impact of climate on 
grass growth and harvest, and on the vegetable crop, plus the direct physical impact of sea ice, 
is discussed first. Second, indirect climatic impact on domestic animals and humans is considered. 

Direct Impacts of Climate 

Grass Growth and Harvest 

During 1816, both grass growth and hay harvesting varied considerably around the country in 
terms of quality and quantity. Only one source, Brandsstadaanndll, from Bléndudalur in 
Huinavatnssysla, gives an unqualified report that the grass grew well. Haymaking began on 25 
July, and there was a successful hay harvest. Another source, that written at Sfdumiuli in 
Myrasysla, categorizes the grass growth as "quite good", and states that the harvest, and 
subsequent use of the hay, also went reasonably well. It is interesting to note that these are the 
only two sources which report a good, or, in the latter case, an "averagely good" summer as far 
as weather is concerned. 

At Ketilsstadir in the east, Sheriff Pall Pérdarson Melsted judged that the grass growth was good 
“on the whole", although "lack of warmth" meant that the outlying pastures did not grow as well 
as the homefields. The harvest, however, was below average. This the Sheriff attributed to the 
damp and rainy summer which prevented the hay drying. The Sheriff of Snefellsnessysla in the 
west, Sigurdur Gudlaugsson, who lived at Gr6f, noted a similar situation. The grass seemed to 
grow well, but in the end turned out to be average. "The harvest from the homefields was very 
mediocre due to rain and damp weather." Writing about the autumn of 1816, he commented 
further: "On account of the autumn’s stormy and inconstant rainy weather, the harvest was very 
poor in many parts of the district, especially on higher ground where some of the hay blew away 
and was washed away from the ground." The opposite situation is reported by Jénas Scheving, 
Sheriff of Borgarfjardarsysla at Leird: "The grass growth, especially in the outlying pastures was 
average, but poorer from the homefields. However, the actual harvesting was excellent." The 
average to poor grass growth he attributed to cold weather from mid-May to about 24 June and, 
more particularly, the dry weather which followed this. The harvest "did not begin until the end 
of this month (July)." The dry weather, which lasted to the end of August, undoubtedly facilitated 
the harvest. The final state of the grass is also said to have been average in the account written 
at Reykjavik. However, grass growth was said to be very late due to cold spring weather. The 
harvest was "difficult." A letter of March 1817, states that in Arnes district, in the south, "the 
weather is supposed to have been not unfavourable to the harvesting of the hay." Furthermore, 
in spite of the difficult harvest, "with the exception of a few individual farms in Kjés district" 
there has not been a lack of hay up to this time. However, the severity of the winter 1816-17 
meant that the upland farmers had to give outside livestock hay almost constantly. "It is thus 
feared that if the winter should remain severe during the present and next month, the lack of this 
item will be considerable." 

In most northern districts, the situation regarding grass and hay during the summer and autumn 
of 1816 seems to have been more difficult than that of most other regions. Stefan Porarinsson, 
writing from Médruvellir in Eyjafjardarsysla, commented that, as a result of the cold spring, the 
grass growth was no more than average in most places in the north. This is echoed by other 
letters from the north. The account from Vidvfk, for example, states that cold, dry weather 
prevented grass growth; and at Grund, Sheriff Gunnlaugur Briem noted that grass growth was 
unfavourable due to a very cold summer. All these northern letters mention that an epidemic, 
which affected people in many parts of Iceland this summer, served to hinder the hay harvest. 
The letter from Vidvfk also commented on the rain that set in during the middle of the harvest. 
Stefan Pérarinsson also noted that long-lasting rain during September, together with storm winds 


that blew some of the hay away, caused a setback to the harvest of the outlying pastures, and 
resulted in this being, in his opinion, below average. 

Comments on grass growth and the harvest in the different sources used here are summarized in 
Table 4. Also included is a summary of the characteristics of the winter, spring and summer 
seasons. In Table 5 the perceptions of the writers on how the weather affected the grass and 
harvest are shown. The main characteristics of the spring and summer seasons, plus grass growth 
and hay yield at each location, are shown in Table 6. Spring weather and grass growth are 
compared, and summer weather and the harvest. There can be little doubt that the summer 
weather directly affected the harvest. For example, if rain or snow or strong winds occurred, the 
harvest would be jeopardized. The exact effect of the spring weather on grass growth is much ~ 
more complex, involving other variables such as soil condition, use of fertilizer, etc., but, from 
previous work (Bergthérsson 1966; Fridriksson 1972; Ogilvie 1981, 1984b; Bergthdérsson et al. 
1988) it is known that unfavourable weather (whether excessively cold, dry or wet) has a 
damaging effect on grass growth. It is interesting therefore to compare the incidence of 
favourable/unfavourable weather with favourable/unfavourable grass growth or harvest in the 
different locations (Table 6). Where these coincide a line is drawn between them. The harvest 
and summer weather agree in every case but one. However, it would be reasonable to assume 
agreement in this latter case also, as the Médruvellir site, where the summer weather was not 
reported, lies only a few kilometres from Grund where the weather was said to be very cold. 
Grass and spring weather, as might be expected, do not agree as well, but the agreement (in six 
out of 11 cases) is nevertheless striking. 

Vegetable Cultivation 

From the latter part of the eighteenth century onwards, a serious attempt was made by the Danish 
authorities, and by enlightened individuals, to get ordinary people to supplement their diet by 
growing vegetables. The most commonly-planted species were potatoes, cabbage and turnips. 
These crops failed almost everywhere in Iceland in 1816. At Vidvfk in the north, for example, 
Sheriff Jon Esp6lfn noted that the number of gardens in use had increased greatly, but that they 
had not done well this year due to "the severe weather and storms" and also to the epidemic 
which affected people almost the whole summer, and prevented them from working. Early in 
1817, he wrote again, commenting that gardening activity had ceased as the ground was frozen. 
He continued: "... one cannot think without sorrow of... the many years of dearth in most places 
in this district..." 

Accounts from elsewhere for 1816 are similar to J6n Esp6lfn’s. Stefan Pérarinsson, writing from 
MO6druvellir, stated that some turnips and cabbage had grown, but that the potato harvest had 
failed completely. Sheriff J6nas Scheving, at Leird, wrote that vegetables had done very badly 
over the past year. This he attributed to lack of sufficient seed, and also to cold spring weather, 
and dry weather in July. A poor vegetable crop also occurred in Vestur-Skaftafellssysla. 

However, the Sheriff there, Lydur Gudmundsson, mainly attributed their "pale and sickly 
appearance" to the effects of the volcanic eruption that occurred under Skaftafellsjékull in June 

As with the grass growth and harvest, it seems reasonable to assume that, aside from the effects 

of this eruption, the weather of 1816 did play a considerable role in the failure of the vegetables. 
This is also suggested by previous work on crop/climatic relationships (e.g., Parry et al. 1988). 


The Impact of Sea Ice 

As noted in the early part of this paper, Iceland is close to the seasonal boundary of Arctic drift 
ice. When the ice reaches Iceland (most commonly, the northern, northwestern and eastern 
coasts) the most striking climatic effect is a lowering of temperatures in the areas affected (see 
also Wilson, this volume, regarding the cooling effect of sea ice lingering near the eastern coast 
of Hudson Bay). Rain and mist may be associated with the ice. The presence of the ice also has 
a direct physical impact. Because the ice prevents access to the open sea or makes it hazardous, 
activities such as fishing and sealing are prevented or hindered. This is no less true today than 
in past centuries, but, in the twentieth century, sea ice has not been common near Iceland. Other 
activities, such as gathering of shellfish from the shore, and the grazing by livestock of seaweed 
and marine plants, are also curtailed by land-fast ice. Such dietary supplements for humans and 
animals are of relatively little importance today, but played a vital role in the past. 

The sea ice did bring some benefits, mainly in the form of driftwood and the occasional beached 
whale or other sea mammal, driven ashore by the encroaching ice. Wood was always in short 
supply, and a whale would greatly augment the food supply. For a more detailed discussion of 
the effects of the sea ice on flora and fauna, see Fridriksson (1969). 

During the period 1809-20, heavy ice years occurred in 1811, 1812 and 1817. During these 
years, ice was present off the northern coasts and elsewhere from some time in January to July 
or August. During 1818 and 1819 very little ice appeared. The former year was very unusual in 
that the sea ice occurred in August, although not for long. In the latter year ice was seen briefly 
in April. 

The year 1816 may be classed as a moderate ice year. During this year, sea ice affected the 
northern coast of Iceland from the beginning of March to the middle of July. Stefan Pérarinsson, 
at M6druvellir, commented that the ice caused persistent northerly winds, frost and cold air. 
Briefly, ice prevented the arrival of the first trading ships at Eyjafjord. At Gardur, in Sudur- 
Pingeyjarsysla, Sheriff Pordur Bjérnsson stated that the seal fishing had been very good until sea 
ice came and prevented this. The shark fishing was poor for the same reason. According to the 
account at Grof, ice also prevented fishing in parts of Breidafjérdur, in the west. But the layers 
of ice "far out to sea" reported by Sigurdur Gudlaugsson, were caused by the sea itself being 
frozen, and not by actual sea ice. The Sheriff commented that in the 11 years he had been there, 
the fishing had never been as poor as this year. 

Although sea ice undoubtedly caused some inconvenience during 1816, there is little evidence to 
suggest that it had a major impact on food supplies. 

Indirect Impacts of Climate 


Most sources mention the severe winter this year, and the frequently frozen ground that prevented 
grazing. Nevertheless, there were no serious losses of livestock. Indeed, only Stefan Pérarinsson, 
writing at M6druvellir in Eyjafjord district in the north, reported that some people lost a number 
of their outside livestock. He wrote: 


Table 4: Summary of Seasons, Grass Growth and Harvest in 1816. 












W - Average 
Sp - Cold, calm 
Sm - Wet 

W - Very severe 
Sp - Sea ice present 
Sm - Rain, snow, frost 

W - Very severe 
Sp - Quite severe 

Smi= 5 2 

Sp ess 

Sm - Very cold 
W - Severe 

Sp - Dry and cold 
Sm - Dry & cold then wet 

W - Mainly severe 

Sp - Weather improved 

Sm - Dry and good during 

W - Fairly severe 

Sp - Variable 
Sm - Variable 
W - Severe 

Sm - Averagely good 

W - Very severe 
Sp - Mainly severe 
Sm - Dry to end Aug. 

W - Severe 

Sp - Unusually cold 

Sm - Mainly cold and 

W - Very severe 
Sp - Severe 
Sm - Unfavourable 


Good on homefields; 
not as good on out- 
lying pastures 




Quite good 

Average on pastures; 
poorer on homefields 



Below average 


Below average 


Very good 


Meagre and spoilt 

' W winter; Sp spring; Sm summer 

Table 5: Contemporary Perceptions of Climatic Impact on Grass Growth and Harvest in 1816. 


Cold spring meant that the outlying pastures did not grow as well as the homefields. Nevertheless, grass 
growth good on the whole. Harvest below average due to wet summer. Not possible to dry hay - therefore 
stacked up damp. 

Harvest poor due to bad weather and epidemic. 


In spite of the cold spring, the grass growth was about or almost average in most places here in the north. 
In the east it is said to have been poorer. The summer’s harvest did not live up to the promise of the grass 
growth, however. This was due to the epidemic which occurred everywhere in the north at the beginning 
of the harvest. Then rains in September plus storm winds adversely affected the hay on outlying pastures. 
Thus, on the whole, harvest below average. 


Grass growth unfavourable due to cold summer. Harvest also, primarily due to epidemic. 

Grass did not grow well due to cold spring and summer weather. Harvest poor due to rains and epidemic. 


In most places the grass growth looked quite good to begin with, but turned out to be only average and, 
on account of wet weather, the harvest of the homefields was mediocre. Due to stormy and inconstant rainy 
weather, harvest very poor in many parts of the district, especially on higher ground where some of the 
hay blew away it was washed away from the ground. 

Dry weather from about 24 June meant that grass growth poorer than last years, so harvest did not begin 

until end July. 

As a result of the cold spring weather, the grass growth was only average and the harvest very difficult. 
Nevertheless, most people do not lack hay. 

Poor grass growth due to volcanic eruption. Harvest spoiled by rains. 



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... the long lasting layers of ice in most places in this region... caused a good many 
farmers, here and there, to suffer a lack of fodder. They therefore lost a number of 
their so-called outside livestock, especially horses, due to emaciation. However, the 
latter loss (of the horses) only applied to some of the inhabitants of Skagafjord and 
Hiuinavatn districts who, to their own detriment, keep far too many horses. On the 
whole, the loss of outside livestock was neither general, nor of great importance. 

Two other letters reported that lack of fodder meant that some livestock had to be slaughtered. 
These letters are from Gardur, in Adaldalur in the north, and Groéf in Snefellsnessysla, in the 
west. The account from the former stated that, although some people were forced to slaughter 
their livestock toward the spring, livestock deaths were not general. The latter source commented 
that in many places people had to slaughter their sheep as the usual winter grass failed in most 

Other sources remark on the difficulties for livestock during 1816, but emphasize that, on the 
whole, they were kept alive. At Ketilsstadir, in the east, the winter was said to be only average 
but, because of large amounts of snow in some places, the outside livestock had to be given 
fodder for a long time. However, "this did not last so long that the animals died of hunger." 
Sheriff Pétur Otteson, writing from Sfdumiuli in the west, stated that, because of the layers of ice 
and snow, virtually no grass was available for the livestock. He continued: "they would have died 
in great numbers if there had not been sufficient fodder after last year’s good harvest. 

The letter from Vik does not comment on the livestock during the winter, but says that, during 
heavy rain and storms from 11 June to 11 July, cows and ewes needed food and shelter. The 
Sheriff added: "After the severe winter, this could scarcely be spared," implying that, here too, 
the livestock needed extra fodder during the winter. According to this letter, the poor hay harvest 
this year caused livestock, especially cows, to be slaughtered in the autumn. From 9 October 
onwards, changeable weather with "hail, snow, layers of ice, rain, sleet and frost" meant that 
little grass was available. The Sheriff at Vfk commented: "The horses and sheep have become 
emaciated and have sometimes needed to be given fodder, and this has had to be shared with the 
few remaining cows." The state of the livestock in the autumn and early winter is also noted by 
the Sheriff of Borgarfjardarsysla, at Leira. After mid-October, "the winter set in with alternating 
frost and drifting snow, thaws and rain. This made it very difficult for the livestock, the horses 
and sheep who need to find their own food, as the frost caused the large quantities of snow and 
water which fell to form a frozen layer on the ground." We may conclude that, during 1816, 
conditions for livestock were, if not easy, not usually difficult either. 

Social Stress 

Research carried out for the period 1601 to 1780 (Ogilvie 1981) has shown that it is very likely 
that during this time climate did play a part in the occurrence of social stress, which manifested 
itself in such phenomena as the desertion of farms, begging and petty crime, plus hunger-related 
diseases and mortality among the people. During 1816, however, such problems were not 
widespread. Only one district reported general difficulties of this kind. This was Snefellsnessysla, 
in the west. Here, Sheriff Sigurdur Gudlaugsson wrote: 

Great lack of food among inhabitants. People pressed by beggars from here and also 
from other districts. The majority of the district’s populace have already got into debt 
at the trading places in previous years, and have scraped together all that they could 
in order to pay. So now they have to give all the best fish to the merchants and have 
little left for themselves except for flatfish and cod’s heads. This is poor winter 


provision, particularly on the coast among the poor fishermen who do not earn 
sufficient during the summer to buy other necessary foodstuffs from the farmers, and 
who therefore frequently live in the greatest misery. 

The lack of food must be partly attributed to the fact that, as the Sheriff noted elsewhere in his 
letter, the trading places were very poorly supplied with corn wares and other imported 
foodstuffs. Furthermore, the fishing, of great importance in this district, largely failed this year. 
Clearly this was largely due to climate. The Sheriff describes how "although there should have 
been fishing in the latter part of the winter months, the severe frost and layers of ice far out to 
sea, frequently prevented the fisherman from getting out to sea for many days on end." 

Because Snefellsnes and nearby areas were important fishing centres, they attracted people whose 

inland sources of food had dwindled. Thus, although most other districts do not report social 
Stresses this year, their silence on such matters may be partly attributable to the fact that the 
people in difficulties had already left to try their luck at the western and southern fishing stations. 


During 1816, most districts in Iceland experienced a very severe winter. One source, the letter 
written by Sheriff Jon Espolfn at Vidvfk, compared it with two other very severe winters in 
recent times, 1784 (see Wood, this volume, regarding climatic effects of the Laki eruption) and 
1802. The spring was also mainly severe in most places. It was a moderate sea-ice year, with ice 
present off the northern coasts from the beginning of March to mid-July. The summer was 
unfavourable, at least for part of the season, in most districts in Iceland with various 
combinations of excessive cold, wet or drought reported. In certain parts, the epithet "year 
without a summer" may have been appropriate, but if we consider the whole summer, over all 
Iceland, then it would not have been. The regional variability reflected in the sources used here 
is quite in accord with what is known of local climatic effects in Iceland (Eythérsson and 
Sigtryggsson 1971; Ogilvie 1984a). 

If the summer of 1816 had been unfavourable in all parts of Iceland, as happened in true "years 
without summers" such as 1756 (Ogilvie 1981) and 1783 (Ogilvie 1986), then the climatic impact 
felt might have been greater. However, it might also have been greater if a favourable harvest 
had not occurred in 1815, thus boosting haystocks. 

It is not difficult to demonstrate that direct impact, for example, on grass growth and hay yield 
did occur in 1816. The indirect role of climate on society this year is harder to define. While it 
is clear that there were difficulties amongst the populace, these were not widespread and were 
compounded by political and economic factors (e.g., by difficulties with trade). Several accounts 
this year report that supplementary foodstuffs received from Denmark were insufficient or of poor 
quality. There were also reports of poor fishing catches. It is true that fish are affected by 
climate, but the relationship is complex and, as yet, not fully documented. Certainly, poor fishing 
catches at sea are not directly linked to climate on land except in the case of heavy storms or 
when lowered temperatures cause ice to form on the sea, thus preventing fishing (as occurred off 
Snefellsnes district this year). The presence of sea ice may also hinder fishing as happened off 
the north coast of Iceland this year. 


In spite of the difficulty in allotting specific roles to economic, political and climatic factors in 
the general well-being of the Icelanders in 1816, there can be little doubt that some indirect 
climatic impact was felt this year. In the climatic context alone, 1816 was certainly an interesting 
year, if not a "year without a summer." 

Dick Harington, Tim Ball and Cynthia Wilson deserve praise for their efforts in organizing the 
meeting " The Year Without a Summer? Climate in 1816" held in Ottawa June 1988. As always, 
I am grateful to many Icelanders for their help. Here I should like to acknowledge in particular 
Pérhallur Vilmundarson, Adalgeir Kristj4nsson and Trausti J6nsson. Part of the research for this 
paper was supported by grant GR3/7013 from the Natural Environment Research Council. This 
paper is dedicated to Valmore C. La Marche Jr. (1937-1988), who had been looking forward to 
joining in the debate on the climate of 1816. 

I had a dream, which was not all a dream 

The bright sun was extinguish’d, and the stars 
Did wander darkling in the eternal space, 
Rayless, and pathless, and the icy Earth 
Swung blind and blackening in the moonless air. 

(From "Darkness" by Lord Byron. Written in 1816.) 


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J6n Espélin, Vidvik. 26 August 1816, no. 2593 and 30 June 1817, no. 2818. 
Sigurdur Gudlaugsson, Grof. 18 February 1817, no. 2811. 

Pétur Ottesen, Sfdumuli. 31 December 1816, no. 2808. 

Johan Carl Thueracht v. Castenskiold, Reykjavik. 17 August 1816, no. 2519 and 5 March 
1817, no. 2847. 

Islands Journal 13 
Pall Pérdarson Melsted, Ketilsstadir. 19 October 1816, no. 392 (formerly Islands Journal 12, 
no. 2645). 

Jénas Scheving, Leir4. 31 July and 31 December 1816, no. 24. 


Lydur Gudmundsson, Vfk. 17 January 1817, no. 27. 

Islenzka Stjérnardeild 8 
Pordur Bjérnsson, Gardur. 23 September 1816, no. 505 (formerly Islands Journal 12, no. 

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530 pp. 


First Essay at Reconstructing the General Atmospheric Circulation 
in 1816 and the Early Nineteenth Century 

H.H. Lamb’ 

Reconstructions of the general atmospheric circulation in January and July year by year back to 
1750, based on the best available network of monthly mean M.S.L. barometric pressure values 
over as much of the world as possible, from observation data in the archives and library of the 
United Kingdom Meteorological Office, were published by Lamb and Johnson (1959, 1961, 
1966). The maps were all analyzed by me, and the analyses were tested by a simulation 
procedure: maps of the years 1919-39 were first analyzed using only restricted networks of data 
corresponding to the information available in the period 1786 to 1820, and these were then 
compared with maps for the same (inter-war) years analyzed with the use of full data. The 
distribution of errors was then studied. On this basis, it was decided that isobars on maps drawn 
for years in the late eighteenth and early nineteenth centuries could be considered satisfactorily 
reliable within regions where the standard error on the test maps was less than 1.0 mb in July 
(or less than 2.5 mb in January, this figure corresponding approximately to the ratio of the 
standard deviation of the observed values in January compared with those in July). 

This meant in practice that isobars could only be presented with confidence over, or very close 
to, Europe between southern Scandinavia, Britain and the western Mediterranean on the maps for 
individual Januarys and Julys in the decade 1810-19. Decade and longer-term mean isobars could 
be reliable over a wider area, spanning most of the Atlantic Ocean between latitudes about 30 and 
50 to 65'N. Isobars at 5-mb intervals were printed as unbroken lines over the areas established 
by the tests as reliable within the limits mentioned (and in regions of slack pressure gradients an 
intermediate isobar might be drawn in at a 2.5-mb interval). 

On the maps for individual Januarys and Julys the isobars were extended, as broken lines, over 
regions where it seemed that the pattern must be broadly reliable, though the pressure values 
could not be relied upon. 

In the case of July 1816 - as with some other seasons of historically dramatic weather - use could 
be made of a wealth of descriptive data on the weather experienced in many places so that it 
seemed reasonable to extend the isobar pattern, as broken lines, far beyond the limits of where 
the pressure values were known. This produced the map for the average conditions prevailing in 
July 1816 (Figure 1). 

The coldness of that summer in eastern Canada, and in northeastern North America generally, 
appears here as attributable to prevalence of air drawn directly from the Canadian Arctic and the 
closeness of a focus of cyclonic activity to Labrador, Newfoundland and off-lying waters. The 
coldness of the summer in Britain, southern Scandinavia and the western part of continental 
Europe is seen to be due to the prevailing concentration of a low pressure region - unusually far 
south for summer - over the areas named, together with indraught of Arctic air from the source 
regions nearby. This is a similar explanation to that more tentatively shown for northeastern 
North America. 

' Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, U.K. 


The much better weather (and crops) experienced in Shetland - and to some extent all over the 
northern half of Scotland - and elsewhere in northern and also eastern Europe, extending south 
to the Crimea, is readily attributable to the higher pressures (and probably greater sunshine) over 
those areas. 


Lamb, H.H. and A.I. Johnson. 1959. Climatic variation and observed changes in the general 
circulation: Parts I and II. Geografiska Annaler 41:94-134. 

. 1961. Climatic variation and observed changes in the general circulation: Part III. 
Geografiska Annaler 43:363-400. 

. 1966. Secular variations of the atmospheric circulation since 1750. Geophysical 
Memoir 110. (Her Majesty’s Stationery Office, for Meteorological Office). London. 
125 pp. 


Weather Patterns over Europe in 1816 

John Kington’ 


An outline of the state of meteorology during the early nineteenth century is presented with 
particular reference to the introduction of the synoptic method of analyzing daily weather maps 
by Heinrich Brandes in 1816. 

Links with the historical weather data made and collected in the 1780s are mentioned in relation 
to the series of daily weather maps for Europe that I am constructing from 1781. 

The feasibility of undertaking a program of similar research for a period of years centred on 1816 
is discussed. As an example, a run of daily charts for Europe in July 1816 is presented together 
with a preliminary analysis of the circulation patterns brought to light in the process. 

Comparisons are made with events in the 1780s, in particular the cold summer of 1784 that 
followed the formation of the exceptional volcanic dust veil after the great eruptions in Iceland 
and Japan the preceding year. 

Historical Weather Data: Comparison of 1816 with the 1780s 

Writing in Breslau, Silesia towards the close of 1816, the German meteorologist Heinrich Brandes 

... If one could collect very accurate meteorological observations, even if only 
for the whole of Europe, it would surely yield very instructive results. If one 
could prepare weather maps of Europe for each of the 365 days of the year, 
then it would be possible to determine, for instance, the boundary of the great 
rain-bearing clouds, which in July [1816] covered the whole of Germany and 
France; it would show whether this limit gradually shifted farther towards the 
north or whether fresh thunderstorms suddenly formed over several degrees 
of longitude and latitude and spread over entire countries ... In order to 
initiate a representation according to this idea, one must have observations 
from 40 to 50 places scattered from the Pyrenees to the Urals. Although this 
would still leave very many points uncertain, yet by this procedure, something 
would be achieved, which up to now is completely new. 

As a meteorological observation network did not then exist, Brandes was unable to examine the 
weather conditions of July 1816 but pursued his hypothesis by making use of data collected 
30 years earlier by the Societas Meteorologica Palatina. Thus the first observations to be studied 
by means of the synoptic method devised by Brandes were those for 6 March 1783 (Figure 1), 
a day on which, like many of those in July 1816, stormy weather prevailed over western and 
central Europe. 

' Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, U.K. 



Carte synoptique 
du 6 Mars 1783. 

PLT | Pe” 
oe RAED 

Figure 1: Synoptic weather map for 6 March 1783 by H.W. Brandes, reconstructed by H. Hildebrandsson. 
Surface wind directions are shown by arrows and the field of pressure by isopleths of equal 
departure of pressure from normal (e.g., -17, -16, -15, etc.). By overcoming the uncertainty 
about the height at which the barometer readings were made, the observations were successfully 
combined to allow the equivalent of isobars to be drawn at a constant level (from Ludlam 1966). 


During the Enlightenment, hopes had been raised in scientific circles that a systematic study of 
meteorological observations would show that the seemingly disordered array of weather variations 
were subject to predictable forms of behaviour. Consequently, extensive networks of observing 
stations were established by two scientific societies in Europe during the early 1780s, namely, 
the Société Royale de Médecine and the Societas Meteorologica Palatina centred at Paris and 
Mannheim, respectively. 

Unfortunately, tangible results proved to be elusive by the statistical approach then applied, which 
earlier had been so successful in predicting the motion of the stars and planets. However, the 
collections of reports from these two societies, together with further data from private individuals 
and ships’ logs, means that a large array of daily instrumental and quantitative observational 
material became, and is still, available for the 1780s over Europe. These data provide a network 
of more than the "40 to 50 places" advocated by Brandes (Figure 2). This information is now 
being subjected to twentieth-century concepts in synoptic meteorology to "yield very instructive 
results" in the construction of daily historical weather maps as envisaged by Brandes 170 years 
ago. As an example of the series of charts now becoming available from 1781, the map for the 
same day as earlier constructed by Brandes (6 March 1783) is illustrated in Figure 3. 

The early blossoming of meteorology in the late eighteenth century was brought to a halt by the 
political confusion and social unrest that followed the outbreak of the French Revolution. The two 
main scientific societies which had been promoting international cooperation in the exchange of 
weather information were disbanded in the mid-1790s. After a lapse of two decades, was it the 
"year without a summer" with its exceptionally cold wet weather and disastrous harvests that 
provided the stimulus for a revival of efforts to understand and predict weather changes? In any 
event, the idea of mapping over a large area simultaneous daily observations of meteorological 
elements such as pressure, wind and temperature (the concept upon which synoptic weather 
studies are based) was, as earlier stated, presented at this time by Brandes. 

Having demonstrated recently that it is indeed possible to map historical weather data on a daily 
basis over Europe for the 1780s (Kington 1988), and knowing that comparable, albeit less well 
sorted and organized, observations are available for 1816, a pilot scheme was initiated for a 
monthly period in that year, following the kind invitation to attend this conference by Dr. C.R. 
Harington. July was chosen for several reasons, not least being the month first highlighted by 
Brandes in his letter of 1816, as quoted above. 

In 1967 the German climatologist, Hans Von Rudloff, examined the weather patterns of 1816 in 
his study of the fluctuations and oscillations of European climate since the beginning of 
instrumental weather observing in the seventeenth century. His analysis showed that there was 
an abnormal distribution of pressure over Europe in the summer of 1816. The subtropical high 
pressure system, the "Azores High", which usually extends northeastwards over the region at 
times during the summer, appears to have been completely absent. Instead, systems of low 
pressure persisting over central Europe allowed polar air streams to be advected farther south 
than normal over the region (Von Rudloff 1967). 

At about the same time as Von Rudloff’s study, Hubert Lamb presented an investigation of 
secular variations in atmospheric circulation since 1750 by means of a series of maps showing 
mean pressure distribution for the months of January and July (Lamb 1967). The chart for July 
1816 (Lamb, this volume) again shows an unusual distribution of pressure, with the "Icelandic 
Low" positioned well to the south of its normal summer latitude. 


O Société Royale de Médecine stations 
A Societas Meteorologica Palatina stations 
@ Private stations 

pis Ship reports (some typical positions) 

Figure 2: Map of stations showing the synoptic coverage available for the 1780s (from Kington 1988) 


Figure 3: Synoptic weather map for 6 March 1783 (from Kington 1988). 


1 6'* MARCH 1400hrs 


Climatic Research Unit 

School of Environmental Sciences 
University of East Anglia 

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NON Sey 7 

Figure 4: Synoptic weather map for 7 July 1816 illustrating the Lamb British Isles Cyclonic weather type. 


a . en 


(2) LP? 

Mi a 

1816 JULY 27 ¢ 

Figure 5: Synoptic weather map for 27 July 1816 illustrating the Lamb British Isles North Westerly 
weather type. 


More recently, in the reconstructions of monthly pressure patterns for Europe back to 1780 based 
on principal components regression techniques, Jones, Wigley and Briffa presented a map of 
pressure anomalies for July 1816 that shows an unusually large negative area, in excess of seven 
millibars, over the British Isles and southern North Sea (Jones et al. 1987). 

All these works strongly indicate that some very pronounced regional anomalies occurred in the 
circulation over Europe in July 1816. Can we discover more? Yes, because an investigation of 
weather patterns on a daily basis can reveal aspects of atmospheric behaviour that are not possible 
to detect from studies made on monthly or longer time scales. 

Although, as previously stated, the two major observation networks of the 1780s were disbanded 
in the following decade, a number of the original stations continued in operation, while others 
were newly established during the early part of the nineteenth century. Using a nucleus of such 
data (readily on hand in the Climatic Research Unit), a run of daily weather maps for July 1816 
was specially prepared for this book. 

The charts have been analyzed and classified with Professor Lamb according to his system of 
British Isles weather types (Lamb 1972). This scheme aims to represent the main types of 
circulation patterns prevalent over the British Isles, namely: Westerly (W), North Westerly (NW), 
Northerly (N), Easterly (E), Southerly (S), Anticyclonic (A) and Cyclonic (C). Since the British 
Isles are centrally placed in the mid-latitude westerly wind belt, as well as being located in one 
of the sectors around the northern hemisphere most frequently affected by blocking of this flow, 
variations in the circulation over the more extensive North Atlantic-European region are also well 
registered by this classification. The classification for July 1816 is given in Table 1. 

A statistical analysis of the classification (Table 2) shows that the circulation over the British Isles 
during July 1816 was strongly dominated by Cyclonic weather types (three times more frequent 
than usual). Of the other patterns, Northwesterly types were also more prevalent than usual (over 
twice as frequent); Southerly types about average; Westerly types, however, were about one third 
of the normal frequency, while Anticyclonic, Northerly and Easterly types were totally absent. 

Typical examples of the two predominant weather types, Cyclonic and Northwesterly, are shown 
in Figures 4 and 5. 

In Table 3 the frequencies of the Lamb British Isles weather types in July 1816 are compared 
with those for 1868-1967, 1781-85 and 1785. 

This shows that frequency values for July 1816 are nearer to the averages for 1781-85 (a period 
in the Little Ice Age) than those of the standard period, 1868-1967. In particular, the circulation 
of 1816 closely parallels that of 1785 with its notable increases in Cyclonic and North Westerly 
types and corresponding decreases in Anticyclonic and Westerly types. 

The Lamb British Isles weather types are also used to determine the PSCM indices of: 
progression, meridionality and cyclonicity, which provide a ready means of indicating the general 
character of the circulation over the region for a period of a month or more (Murray and Lewis 


Table 1: Lamb British Isles Weather Types for July 1816.' 

1 (e, 11 CNW 21 Cc 

Z € 12 NW 22 SW 

3 (@ 13 WwW 23 CS 

4 c 14 C 24 Cc 

5 (© 15 Cc 25 ( 

6 Cc 16 C 26 WwW 

7 C ity € 27 NW 

8 € 18 Cc 28 NW 

9 Cc 19 C 29 CNW 
10 Cc 20 S 30 NW 

Sih U 

' C Cyclonic, CNW Cyclonic North Westerly, NW North Westerly, S Southerly, SW South Westerly, CS Cyclonic 
Southerly, W Westerly, U Unclassified. 

Table 2: Lamb British Isles Weather Types. Monthly Frequencies for July 1816 with Long-Period Mean 
Percentage Values Given in Brackets for Comparison. ' 

Days % % 
Ww 141 2" 8 (26) 
NW v7 oO hee 5) 16 (7) 
N fs 0 0 (7) 
E gan 0 0 (4) 
S | Wey Wp) 2 6 (5) 
A wits 0 0 (24) 
(C UT) Ty ly 20'% 66 (22) 
U 1 1 3 (5) 

' W Westerly, NW North Westerly, N Northerly, E Easterly, S Southerly, A Anticyclonic, C Cyclonic, U 

In July 1816: P=-12 or -3; S=-3 or -1; C=+41 or +42 and M= 11 or 13’ 
Thatas: Py Sj, Cy Me 

This shows that the circulation over the British Isles in July 1816 was characterized by blocked 
or quasi-stationary cyclonic weather systems. The C index value of +41 or +42 is far greater 
than the maximum value of +30 (1936) in the official long-period record from 1861. 
Interestingly this record was also broken in the 1780s when the C index in July 1785 was +33. 

In July 1785: P=-3; S=-14; C=+33; and M=14 
That is: P5 Ss; €, M, 

' The slightly differing results are due to dealing with a run of charts from a single isolated month, resulting in a 
certain lack of synoptic continuity at the beginning and end of the series. 


ws \ Ape 
SNe —— 
2 oe 



Sa te 
JULY 1816 

Figure 6: Rainfall anomalies (%) for July 1816. 

Table 3: Lamb British Isles Weather Types. July Frequencies for 1816 and 1785; Period Average 
Frequencies for 1868-1967 and 1781-85.' 

Number of Days 

Ww NW 
1816 2 5.0 
1868-1967 8.1 Zee 
1781-85 6.9 a9 
1785 1.0 6.5 








' W Westerly, NW North Westerly, N Northerly, E Easterly, S Southerly, A Anticyclonic, C Cyclonic. 

Thus there is a striking similarity (blocked and very cyclonic) in the PSCM "signatures" of 1816 

and 1785. 

As rainfall over England and Wales has been found to be closely correlated with the C-index, it 
is not surprising that very heavy falls of rain occurred over the region in July 1816 (Figure 6). 
The map, however, shows that it was not uniformly wet over the British Isles or continental 
Europe. For instance, while rainfall over southwestern Ireland, southern Wales, southwestern 
England, most of France, parts of Belgium, Holland and western Germany exceeded 200% of 
normal, northwestern Scotland, Orkney, Shetland, Denmark, Norway and Italy were drier than 
usual. Contemporary accounts confirm this contrasting pattern of wet and dry regions: 


Melancholy accounts have been received from all parts of the Continent of the 
unusual wetness of the season; property in consequence swept away by 
inundation, and irretrievable injuries done to the vine yards and corn crops. 
In several provinces of Holland, the rich grass lands are all under water, and 
scarcity and high prices are naturally apprehended and dreaded. In France, the 
interior of the country has suffered greatly from the floods and heavy rains. 


"The Norfolk Chronicle", 20 July 1816 

With depressions centred over or near Ireland for most of the month, the weather over the 
country was very unsettled and wet. Apparently, all parts had more rain than usual, with the 

extreme southwest probably having more than twice the normal amount (Figure 6). 

The summer and autumn were excessively wet and cloudy. ... the sun was in 
general obscured by clouds during the months of July, August and September. 

Great thunderstorms occurred during the month of July, accompanied with 
hail of an unusually large size. These storms were general throughout the 



July - wet, great storms, and inundations in England and Scotland, as well as 
throughout this country ... The month was, without, perhaps, the exception 
of a single day, a continuity of showers of hail or rain, and at the same time 
very cold. 

Snow remained on some of the hills in Scotland until the middle of July, 
during which month great thunderstorms occurred in England. 

In consequence of the incessant rain, there is a great blight in the wheat crop, 
particulary in Wicklow and Tipperary: the rain was so severe that scarcely 
any corn was left standing. For many years so untoward a season had not 
been experienced, not one week of fine weather since May. Eight weeks of 
rain in succession. Hay and corn crops in a deplorable state. The grains of 
corn in many places are covered with a reddish powder like rust, which has 
proved very destructive to the crop, especially in the counties of Kilkenny and 
Antrim.’ The wheat crop was especially injured. Great floods occurred in the 

The fields of corn presented a lamentable appearance, in many places being 
quite black. Before the crop was reaped, re-vegetation had commenced, and 
green shoots were perceived on the fields. 

The harvest of grain was uncommonly late both in this country and in 
England; corn remained uncut during the latter parts of October and 
November, and much of it was altogether lost. The cold of this season proved 
highly injurious to the crop of potatoes also. These, which constitute the 
principal or only food of the poor in most parts of the country, were small 
and wet, and probably more defective in nutriment than the grain. 

The potato crop both in England and Scotland was defective. 
"The Census of Ireland," 1851 

This month for the most part good weather. Quite warm 21° and frequent 
rain, although this did not do any harm. On the other hand in Germany and 
Switzerland terrible damage occurred with rivers flooding. This was caused 
by persistent rain ... whole tracts of land were under water. The hay harvest 
was also ruined in England. 

"A Jutland Weather Diary" (Ribe) 

' This may have been the result of volcanic aerosol particles being washed out of the atmosphere by the rain which, 
in turn, might have been intensified by the increase in condensation nuclei. Editor’s note: Perhaps the possibility 
of fungal rust should be considered also. 


Western Russia and the Baltic Sea Coast 
The city of St. Petersburg [Leningrad] has for a month past suffered by 
drought and prayers for rain have been offered up at Riga and Dantzig while 
Germany is devastated by inundations and the churches of Paris are filled with 
suppliants praying the Almighty for dry weather. 

"Records of the Seasons" 


One of the main objectives of this book has been to determine to what extent the Tambora | 
eruption in 1815 affected world climate. Already we know that some mid-latitude regions of the 
northern hemisphere, such as eastern North America and western Europe, were much cooler than 
normal in the following year, 1816. There is an interesting parallel in the 1780s when it is 
estimated that annual mean temperatures in mid-latitudes fell by 1.3°C after eruptions in Iceland 
and Japan in 1783. However, there appear to be two major points of difference: the timing and 
length of cooling. By all accounts it appears that, unlike 1815-16, the cooling signal in the mid- 
1780s was strongest not in the year immediately following the eruption but in 1785, two years 
after the event. Nevertheless I have shown that there were some notable similarities in the 
circulation patterns of the two cold years, 1816 and 1785. Furthermore, the marked increase in 
cyclonicity over the British Isles in July 1816 is in accordance with Lamb’s (1977) finding that 
there is a tendency for the subpolar low-pressure zone (the "Icelandic Low") to be displaced 
southwards over the British Isles during the first July after a great eruption, resulting typically 
in cold wet summers over the region. Another area of cyclonic activity near Newfoundland gave 
similar weather conditions over eastern North America. However, the volcanic signal apparently 
soon died away, with temperatures recovering to above normal values by 1818. On the other 
hand, after high-latitude eruptions (e.g., those of 1783), pressure and related temperature 
anomalies in mid-latitudes appear to persist longer - the circulation patterns determined for July 
in the cold year of 1785 confirm this trend. 


Drs. A.E.J. Ogilvie and P.D. Jones kindly helped in processing various historical weather data 
from the archives of the Climatic Research Unit. Observations from Dublin and France were 
kindly supplied by Dr. J.G. Tyrrell (University College, Cork) and Dr. D. Hubert (Observatoire 
de Meudon), respectively. 


Baker, T.H. 1883. Records of the Seasons, Prices of Agricultural Produce and Phenomena 
Observed in the British Isles. Simpkin, Marshall and Co., London. 

Dublin. 1856. The Census of Ireland for the Year 1851. H.M.S.O., London. 
Jones, P.D., T.M.L. Wigley and K.R. Briffa. 1987. Monthly mean pressure reconstructions for 
Europe (back to 1780) and North America (to 1858). DOE Technical Report No. 37, 

United States Department of Energy, Carbon Dioxide Research Division, Washington, 


Kington, J. 1988. The Weather of the 1780s Over Europe. Cambridge University Press, 

Lamb, H.H. 1972. British Isles weather types and a register of the daily sequence of circulation 
patterns, 1861-1971. Geophysical Memoirs No. 116. H.M.S.O., London. 

. 1977. Climate: Present, Past and Future, Volume 2, Climatic History and the Future. 
Methuen, London. 

Ludlam, F.H. 1966. The Cyclone Problem: A History of Models of the Cyclonic Storm. Imperial 
College of Science and Technology, London. 

Murray, R. and R.P.W. Lewis. 1966. Some aspects of the synoptic climatology of the British 
Isles as measured by simple indices. Meteorological Magazine 95:193-203. 

Von Rudloff, H. 1967. Die Schwankungen und Pendelungen des Klimas in Europa seit dem 
Beginn der regelmdssigen Istrumenten-Beobachtungen (1670). Vieweg, Braunschweig. 


The Climate of Europe during the 1810s with Special 
Reference to 1816 

K.R. Briffa! and P.D. Jones! 

The long climatic records available for Europe are used to place the seasonal temperature, 
precipitation and sea-level pressure anomaly maps for 1816 into their longer-term context. The 
prevailing climate of the decade of the 1810s (1810-19) is also described with reference to 
modern climatic normals. The 1810s were probably one of the coldest decades recorded over © 
Europe since comparable records began about 1750. It was only the weather during the spring 
and, more particularly, the summer of 1816 that was highly anomalous with respect to both 
recent normals and those for the 1810s. 

Tree-ring-based reconstructions of temperature for a ‘summer’ (April-September) season are 
available in the form of anomaly maps back to 1750. They indicate that the summer of 1816 was 
the coldest since 1750 in Britain, that it was the second coldest (after 1814) in central Europe and 
that in Scandinavia conditions were near normal. 


Many studies have considered the weather extremes that occurred during the summer of 1816, 
the so-called "year without a summer" (Landsberg and Albert 1974; Stommel and Stommel 
1979). Studies have tended to concentrate on the particular season itself, rather than considering 
the weather and climate of the rest of 1816 and the decade of the 1810s. 

In this article we propose to make use of the long records of temperature, precipitation and mean 
sea-level pressure (MSLP) available for most of Europe. We will describe seasonal anomaly maps 
for 1816 with respect to twentieth century reference periods and in relation to those of the 1810s 
(defined here as 1810-19). We also compare the climate of the 1810s to recent reference periods. 

Finally, previously published maps of mean April-September temperature reconstructed from a 
network of maximum-latewood-density tree-ring chronologies in Europe are reproduced for each 
of the years 1810-19. 


Instrumental recording of air temperature and precipitation totals extends back in Europe to the 
late seventeenth century. Most of the pre-twentieth century data have been assembled in computer 
compatible form in data archives. Here we use the compilation of air temperature and 
precipitation data produced by Bradley er al. (1985). This archive contains temperature data for 
46 stations in Europe with series that extend over most of the years of the 1810s (Table 1; 
Figure 1). Of these 46 stations, 12 do not have comparable data through to and encompassing the 
twentieth century. We can still use these more restricted data, however, to compare the average 
temperature of 1816 to that of the 1810s. 

! Climatic Research Unit, University of East Anglia, Norwich NR4 7TJ, U.K. 


Table 1: Names and Locations of Stations with Temperature Data for the 1810s 
Continuous to the Present Day. 

Lat.(°N) Long. 

1. Trondheim 64.3 10.5E 

2. Stockholm 59.4 18.1E 

3. Torneo 66.4 23.8E 

4. Woro 63.2 22.0E 

5. Gordon Castle 57.6 3.1W 

6. Edinburgh 35-9 3.2W 

7. Manchester 53.4 2.3W 

8. Greenwich SES 0) 

9. Copenhagen S557 12.6E 
10. De Bilt Soe 3:26 
11. Basel 47.6 7.6E 
12. Geneva 46.2 6.2E 
13. Montdidier 49.7 2.6E 
14. Chalons 48.9 4.4E 
15. Paris 48.8 D235) 8 
16. Strasbourg! 48.6 7.6E 
17. Nice 43.7 IQE 
18. Berlin 3225 13.4E 
19. Karlsruhe! 49.0 8.4E 
20. Stuttgart! 48.8 9.2E 
21. Regensberg' 49.0 1218 
22. Augsburg’ 48.4 10.4E 
23. Munchen! 48.1 11.7E 
24. Hohenpeissenberg' 47.8 11.0E 
25. Kremuenster! 48.1 14.1E 
26. Wien Hohe Warte 48.2 16.4E 
27. Innsbruck 47.3 11.4E 
28. Klagenfurt 46.7 14.3E 
29. Prague 50.1 14.3E 
30. Leobschutz 50.2 17.8E 
31. Gdansk 54.4 18.6E 
32. Warsaw SA) 21.0E 
33. Wroclaw S151 17.0E 
34. Budapest 47.5 19.0E 
35. Udine 46.0 13.1E 
36. Turin 45.2 7.7E 
37. Milan 45.4 9.2E 
38. Padua 45.4 12.0E 
39. Bologna 44.5 11.5E 
40. Rome 41.7 12.5E 
41. Palermo 38.1 13.4E 
42. Arkhangel 64.6 40.6E 
43. Leningrad 60.0 30.3E 
44. Vilnjus 54.6 25-5E 
45. Kazan Seite 49.1E 
46. Kiev 50.5 30.5E 

' Not labelled on Figure 1. 


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Locations of the 47 sites with instrumental temperature data during the 1810s. Details of the 

sites are given in Table 1 (from Bradley et al. 1985). 

Figure 1: 

Although the Bradley ef a/. (1985) compilation contains details of the sources of the data and the 
methods, where known, by which the observations were made, it does not consider the long-term 
homogeneity of the individual station data sets. The homogeneity of the station temperature data 
used here has, however, been assessed by Jones et al. (1985). 

For precipitation, the Bradley et a/. (1985) compilation contains data for 29 sites covering the 
1810s. Assessment of the homogeneity of the precipitation data is a considerably more difficult 
task than for air temperature. Although the stations used here have not been assessed for 
homogeneity, the data for 27 of the sites are among the 180 or so homogeneous European 
precipitation records assembled by Tabony (1980, 1981). Only the data for Warsaw and Prague 
are not in this set. The locations of the 29 precipitation sites we have used are shown (Figure 2; 
Table 2). 

Some of these early temperature and precipitation series have been used in conjunction with early 
station pressure records by Jones et al. (1987) to reconstruct gridded monthly-mean mean sea- 
level pressure values (MSLP) over Europe extending back to 1780. Jones et al. used a principal 
components regression technique that involves fitting equations expressing MSLP at individual 
grid points in terms of pressure, temperature and precipitation series at all stations in the 
predictor network. The fitting was carried out over a 75-year (1900-74) ‘calibration’ period, and 
the reliability of the gridded reconstructions was assessed by comparing the estimated data with 
actual observations over an independent ‘verification’ period, 1873-99. Jones et al. (1987) 
showed that over Europe, between 65-40°N and 10°W-30°E, the reconstructions are of high 
quality with 80% or more of the variance of the observed pressure data being explained in each 
of the separate monthly reconstructions. 

From this bank of reconstructed pressures we have extracted the data for individual months and 
averaged them to produce maps of MSLP anomalies for the four standard seasons of the year 

Anomaly Maps 

Figure 3 shows seasonal temperature anomaly maps for 1816 with respect to the reference period 
1951-70. Winter in this and subsequent figures is taken to be December 1815 to February 1816. 
All four seasons are shown to have been generally cooler in 1816 compared with the reference 
period. Warmer conditions were experienced only over northern Mediterranean coasts and 
European parts of the Soviet Union, and then only in spring, summer and autumn. The most 
anomalously cool regions were Scandinavia and northern British Isles (during winter, spring and 
autumn) and central Europe (in summer). 

Figure 4 shows seasonal precipitation anomaly maps for 1816 (expressed as percentages of the 
1921-60 reference period). Most regions of western Europe were drier than normal except for 
summer. Below normal precipitation is evident over central and southern Europe in winter and 
to some extent in spring and autumn. During summer the only relatively dry areas were southern 
Italy and northwestern Scotland. 


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Figure 2: Location of the 29 sites with precipitation-gauge data during the 1810s. Details of the sites are 

given in Table 2 (from Bradley et al. 1985). 


Table 2: Names and Locations of Stations with Precipitation Data for the 1810s Continuous 
to the Present Day. 

Lat.(°N) Long. 
1. Uppsala 59.9 17.6E 
2. Lund S5)57/ 13.2E 
3. Inverness ie 4.2W 
4. Eallabus 55.6 6.2W 
5. Edinburgh Spe) 3.2W 
6. Manchester 53.4 2.3W 
7. Mansfield SB iol 1.1W 
8. Podehole 52.8 0.1W 
9. Kew S125 0.3W 
10. Oxford Se7/ 1.2W 
11. Hoofdoorp S203 4.7E 
12. Lille 50.6 3.1E 
13. Montdidier 49.7 2.6E 
14. Paris 48.8 2.5E 
15. Nancy 48.7 6.2E 
16. Strasbourg 48.6 7.6E 
17. La Rochelle 46.1 1.1W 
18. Toulouse 43.6 1.4E 
19. Marseille 43.3 5.4E 
20. Trier 49.8 6.7E 
21. Karlsruhe 49.0 8.4E 
22. Klagenfurt 46.7 14.3E 
23. Prague 50.1 14.3E 
24. Warsaw 52.2 21.0E 
25. Udine 46.0 13.1E 
26. Milan 45.4 9.2E 
27. Padua 45.4 12.0E 
28. Bologna 44.5 11.5E 
29. Rome 41.7 12.5E 

In Figures 5-8 we show similar seasonal anomaly maps for temperature and precipitation, placing 
the 1810s in the context of modern reference periods, and 1816 in relation to the 1810s. The 
1810s (Figure 5) were colder than the recent reference period during winter and autumn but were 
somewhat milder during spring and summer, particularly over eastern and southern Europe. The 
relative coolness of the 1810s with respect to 1951- 70 in winter and spring means that 1816 was 
less anomalous when viewed against this decade as a whole (Figure 7). For precipitation, the 
1810s were generally drier than the 1921-60 reference period (Figure 6), other than over the 
British Isles and Scandinavia in spring and Italy during summer and autumn. 


wo (2) wo oO wo [o} wo 
ive) w + + mM 
ious Lima Salome re 



1816 w.r.t. 1951-70 

° il 
7 \ 7 = 
io co) 2c = 
10) a | I 2 Wy 2 ° 
ve) TT) wl es fe 
= Ss ~ = >) o 
egy s ba Q 
i ® =>) 
ci = fet ‘ =) {= 
a > © £ 
o 2 WATS r 
= 2 2 
= | 
walt L | eile 
Oo ‘9) fo) w fo) w fo} wo [o} wo fo) V9) oO wo 
6 o o w wo + + o Te) wo wo + + a) 

Figure 3: Seasonal temperature anomaly maps in degrees Celsius for 1816 with respect to the 1951-70 
reference period. Winter is the average for December 1815 to February 1816. Spring 
(March-May), summer (June-August) and autumn (September-November). 








O9-L26l “V4M OLB8L 

O9-Lc6L ¥4™M QIBL 


sg uoljyeyIdi9edg 

O9-Lc6l }4M QILBL 

O9-LZ6l }4M OIBL 

OS Ov o¢ 10Y4 Ol 0 Ol- Ov o¢ 0z OL ) ORS 

Seasonal precipitation departures for 1816: values expressed as percentages of the 1921-60 

reference period mean. 

Figure 4 















0) Ol- 


OZ-1S6l ¥4™ SOLEL 


0) (Ol 

Seasonal temperature anomaly maps for the 1810s (1810-19) with respect to the 1951-70 

reference period. 

Figure 5 


09-1261 Y4™M SOLEL 


O9-Lc6l Y4M SOLBL 

09-1261 }4M SOLBL 


Seasonal precipitation departures for the 1810s: values expressed as percentages of the 1921-60 

reference period mean. 

Figure 6 


oO o 
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e = 2 
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SANS ! 1 
O° 9) ° 
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[o) w w jo} wo Q ve) fo) w oO w 
U t+ w+ ve) w wo + t+ Le) 

Figure 7: Seasonal temperature anomaly maps for 1816 with respect to the average for the 1810s 


uOIeydI9edq uolejyIdioadg 
SOL8L }4M OLBL : , SOl8L “}4M QIBL 

St é .A Sr Sr 
0oL 001G > : 
001 stem 
Ss : ya Ss ss 
09 Re 09 09 
s9 UoNeydI9edd so HoleNdIoe el s9 
ita go SOLBL “V4 SLeL ee SOL8L ‘YM Lgl 
OL ie el 242 [ “a, st OL 
os Ov o¢ oz OL 0 O1- Ov o¢ 0z ol 0 OL- 

Seasonal precipitation departures for 1816: values expressed as percentages for the mean value 

for the 1810s. 

Figure 8 

Comparison of the 1816 Seasonal Temperature, Precipitation and MSLP Anomaly Maps | 

Winter (Figure 9) 

During the winter of 1815-16 all of Europe was affected by anomalously low pressure with 
respect to the 1941-70 reference period. Both southern and western Europe were affected by 
greater advection from eastern Europe, which would tend to bring drier and cooler conditions to 
these regions. Enhanced northerly circulation over western Europe resulted in below normal 
precipitation in all areas except north-facing coasts. 

Spring (Figure 10) 

Europe is again shown to be almost entirely under the influence of lower-than-normal pressure. 
The negative pressure anomaly in this season is centred over northern France, implying that both 
Britain and Scandinavia experienced increased easterly and northeasterly weather, from Finland 
and the Gulf of Bothnia, leading to cold temperatures. Drier conditions over continental Europe 
may have resulted from the stagnation of a number of depressions over this region. Wetter than 
normal weather over northern Europe was associated with a greater degree of air flow over 
adjacent seas. 

Summer (Figure 11) 

Virtually the whole of Europe was affected by anomalously low pressure centred over northern 
Germany and Denmark. Milder conditions prevailed over European parts of the Soviet Union 
because of the influence of increased southerly flow across these areas. Britain and the rest of 
western Europe were affected by anomalous northerly and northwesterly airflow bringing cooler 
temperatures. The coldest conditions of the summer occurred in northern Alpine regions. Over 
Scandinavia, in contrast to winter and spring, conditions were near to the recent normal. 
Precipitation was considerably greater over northern France and southern England. 

The summer (June-August) of 1816 was the coldest recorded in the Central England temperature 
series (Manley 1974; updated in Jones 1987). Temperatures were 2.2°C colder than the 1931-60 
average. The Manley series extends back to 1659 (though with slightly lower reliability before 

Autumn (Figure 12) 

Again most of Europe was under the influence of anomalously low pressure, although less intense 
than in the other seasons. The centre of the anomaly was located over Poland, whereas pressure 
was near normal over Ireland and Scotland. Enhanced northerly and northeasterly air circulation 
over Scandinavia, Britain and central Europe, particularly north of the Alps would have led to 
cooler than normal conditions in all of these regions. Southern Europe and European parts of the 
Soviet Union were milder as a result of anomalous westerly and southerly airflow, respectively. 

Precipitation anomalies are consistent with the circulation patterns. Reduced westerly airflow 
would give drier conditions in northern Germany and southern France. Enhanced north and 
northeasterly flow is consistent with above average precipitation over northern coastal Sweden 
and some southern North Sea coasts. 

Inferences from Tree-Ring Parameters 

Beside the long instrumental-based climatic records available for Europe back to the early 1800s, 
detailed year-by-year maps of ‘summer’ (April-September) temperature across Europe have also 


° ° 
KR o 
a J c 
re} © i ° 
Oas Oo) 
-u= ~~ Wo 
res :E# 
ve = or Q 
“Ze 2 ie 
SESE 223 
© a 
te g 
roe} o 
i - 

WINTER 1816 

Figure 9: Climate anomaly maps: Winter. Mean sea level pressure anomalies with respect to 1900-74, 
air temperature anomalies with respect to 1951-70, precipitation as percentages of the 1921-60 

reference period. 


(Si w 

™ ive) 

fo) ~ 

wo \ T 


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8 \ 
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ro) SS 
= \ 
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SPRING 1816 

Figure 10: Climate anomaly maps: spring. 






1816 w.r.t. 1921-60 



65 B 





















09-1261 “34M 918L 


9181 HAWWNS 

Climate anomaly maps: summer. 

Figure 11 



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e Ne! 
ie 2 we aunjesodway | 
69) - ——a\ SSS ot NWALAY Ad so 
RS eSSond =< O19 OLZ-LS6L ‘¥4'™ QLBlL 
r S =—— 
ang _ 

OL Sai coma OL 

OS Ov O¢ 0Y4 Ol 0) (O}|s—= 


Figure 12: Climate anomaly maps: autumn. 





i) ) 




(April to September) temperature based on maximum latewood 

Reconstructions of ‘summer’ 


Figure 13 

density chronologies from trees at 37 high-altitude and high-latitude sites across Europe. 
Further details of the reconstructions are in Briffa et al. (1988) and Schweingruber et al. 



been produced from a network of tree-ring chronologies (Briffa et al. 1988; Schweingruber et» 
al. 1989). These maximum latewood density chronologies, 37 in all from trees at high altitude 

or high latitude, provide reliable reconstructions of summer half-year temperature for the area 

between 45 to 70°N and 0 to 30°E (i.e. most of western Europe north of the Mediterranean) 

back to 1750. There are instrumental data for only a few stations (in London and central Europe) 

back as far as 1750. 

The dendroclimatic reconstructions show that the summer of 1816 was cold, particularly in the 
United Kingdom and central Europe. They also show that summer in Scandinavia was not 
unusual. These results agree with instrumental temperature data. In the United Kingdom, the 
relative cold of the summer of 1816 was unmatched in any other year from 1750 to the present, 
and in central Europe there has been only one colder year - 1814. 

All of the reconstructed maps, plotted as anomalies in degrees celsius from the 1951-70 reference 
period mean, have been published by Schweingruber et al. (1989). In Figure 13 we reproduce 
the maps for each of the ‘summers’ from 1810 to 1819. It is clear from these maps that cool 
conditions prevailed over almost all of Europe during the summers of 1812, 1814, 1816, 1817 
and, to a lesser but still significant extent, 1813. The magnitudes of the negative temperature 
departures were however clearly greatest in 1816, and the most extreme of these (temperature 
departures below -2°C) were experienced in southern Britain and northern France. 


The climate and weather of 1816 was indeed unusual over Europe, particularly during the 
summer. The evidence indicates generally cool conditions over the whole of western Europe 
during all four seasons of this year. The cold was particularly marked in summer when it was 
also wetter everywhere in Europe except in the eastern Mediterranean. Temperature and 
precipitation maps for the four seasons of 1816 are consistent with the reconstructed circulation 
maps. When considered in the context of the decade of the 1810s however, only the weather 
during the summer can be classed as extreme. The 1810s were an exceptionally cool/dry decade 
compared with modern reference periods. 

Tree-ring reconstructions also show the ‘summer’ (April to September) of 1816 to have been 
cool, most particularly over Britain. These reconstructions show that summers from 1810 to 1819 
over much of western Europe were generally cool. However, as with the instrumental data, the 
tree-ring reconstructions indicate that summers as cool as that of 1816 also occurred in each of 
the years from 1812 to 1814, i.e. before the eruption of Tambora. 


The authors are grateful to Dr. F.H. Schweingruber for Figure 13. 


Bradley, R.S., P.M. Kelly, P.D. Jones, C.M. Goodess and H.F. Diaz. 1985. A climatic data 
bank for northern hemisphere land areas, 1851 to 1980. United States Department of 

Energy Technical Report TRO17, United States Department of Energy, Carbon Dioxide 
Research Division. Washington, D.C. 335 pp. 


Briffa, K.R., P.D. Jones, T.M.L. Wigley, J.R. Pilcher and M.G.L. Baillie. 1986. Climate 
reconstruction from tree rings: Part 2, Spatial reconstruction of summer mean sea-level 
pressure patterns over Great Britain. Journal of Climatology 6:1-15. 

Briffa, K.R., P.D. Jones and F.H. Schweingruber. 1988. Summer temperature patterns over 
Europe: a reconstruction from 1750 A.D. based on maximum latewood density indices of 
conifers. Quaternary Research 30:36-52. 

Jones, D.E. 1987. Daily Central England temperature: recently constructed series. 
Weather 42:130-133. 

Jones, P.D., S.C.B. Raper, B.D. Santer, B.S.G. Cherry, C.M. Goodess, R.S. Bradley, 
H.F. Diaz, P.M. Kelly and T.M.L. Wigley. 1985. A grid point surface air temperature 
data set for the northern hemisphere, 1851-1984. United States Department of Energy 
Technical Report TRO22. United States Department of Energy, Carbon Dioxide Research 
Division. Washington, D.C. 251 pp. 

Jones, P.D., T.M.L. Wigley and K.R. Briffa. 1987. Monthly mean pressure reconstructions for 
Europe (back to 1780) and North America (to 1858). United States Department of Energy 
Technical Report TRO37. United States Department of Energy, Carbon Dioxide Research 
Division. Washington, D.C. 99 pp. 

Landsberg, H.E. and J.M. Albert. 1974. The summer of 1816 and volcanism. Weatherwise 

Manley, G. 1974. Central England temperatures: monthly means 1659-1973. Quarterly Journal 
of the Royal Meteorological Society 100:389-405. 

Schweingruber, F.H., K.R. Briffa and P.D. Jones. 1989. (in preparation). 

Stommel, H. and E. Stommel. 1979. The year without a summer. Scientific American 

Tabony, R.C. 1980. A set of homogeneous European rainfall series Met 0.13 Branch 
Memorandum No. 104. United Kingdom Meteorological Office, Bracknell. 

. 1981. A principal component and spectral analysis of European rainfall. Journal of 
Climatology 1:283-291. 


The 1810s in the Baltic Region, 1816 in Particular: Air Temperatures, 
Grain Supply and Mortality 

J. Neumann! 


The mean acidity of the ice core from Créte, central Greenland, for the layer dating to 1816, one 
year after Tambora’s eruption, has been found by Hammer et al. (1980) to be nearly three times 
greater than that of the layer dating to 1884, one year after Krakatau’s eruption. Despite the © 
aforementioned fact, air-temperature data of the Baltic meteorological stations that took 
observations both in the 1810s and the 1880s (Copenhagen, Gothenburg, Stockholm, Trondheim 
and Uppsala), do not show that the coldness of 1816 relative to 1814 was any greater than that 
of 1884 relative to 1882. Moreover, the year 1812 was much colder than 1816 when the two are 
compared with 1814 at all Baltic stations, although no known important eruption took place 
shortly before 1812. It seems plausible that the plumes reaching the Baltic Region following the 
two eruptions were too ‘thin’ to have produced any appreciable effect on air temperatures. 

An examination of data on grain harvests in Denmark, Finland, Norway and Sweden does not 
indicate that either in 1816 or 1817 there was any noteworthy crop failure. In contrast, the year 
1812 (a cold year) was marked by shortfall of the harvest, in consequence of which in 1813 there 
was a partial famine in Norway, partly because of war conditions (blockade by the British Navy) 
it was hard to get supplies from abroad. 

Mortality data are also available for the above four countries. Mortality was relatively high in 
1812 and/or 1813, but not in 1816-17. 

No harvest or mortality data are available for Russia. Lists of famines in Russia show none in 
1816. In 1817 there was a price rise in a limited area of the Empire. 

All-in-all, the Baltic Region did not suffer from Tambora’s eruption unlike the lower 
mid-latitudes of western and central Europe. It is suggested that the Baltic Region, as well as 
southern European Russia, were spared as they were crossed by air masses whose stratosphere 
had become depopulated of small volcanic particles, while the troposphere became cleansed of 
particles through washout by rain previously. 


Did the particle cloud originating and developing (Appendix 2, No. 1) from Tambora’s eruption 
in 1815 reach the relatively high latitudes of the Baltic Region? Was the year 1816 without a 
summer in the region? Was there, perhaps a famine in either 1816 or 1817, or both, as in many 
areas of western, southwestern and central Europe and North America? 

| Emeritus, Department of Atmospheric Sciences, The Hebrew University, Jerusalem, Israel. In 1986-90 visiting 

with the Department of Meteorology, University of Copenhagen, Denmark. 


Hammer et al. (1980, Figure 1) have shown that the acidity of the ice layer dating to 1816 at the 
Créte site in central Greenland (71°N) amounts to 5u equiv. H* per kg ice. For comparison, the 
acidity of the layer dating to 1884, the year following Krakatau’s eruption, amounts only to 2.5 
(here and below, same units as above). If we subtract from these figures the background activity 
of 1.2 + 0.1, then the contribution of 1816 works out nearly three times as great as that of 1884. 
Observations for 1816 are consistent with the assumption that aerosols of the Tambora eruption 
did reach a high northern latitude, at least on the western side of the North Atlantic. 

The acidity accretion of 3.8 + 0.1 equiv. H* per kg ice in 1816 really is a small mass for a 
full year’s deposition. Only the sensitive physical method applied by Hammer ef al. (1980) 
allowed its measurement to such a high degree of precision. We do not know how to turn this 
deposition into figures representing an atmospheric plume of particles. Even if it turned out that 
the plume reaching Greenland extended over an air layer of appreciable thickness and particle 
concentration, there is no guarantee that the plume reaching the Baltic Region was similarly 
‘powerful’ for scattering back to space a significant fraction of direct solar radiation. 

In order to examine the possible effects of Tambora’s eruption on the Baltic Region, we shall 
consider the available air-temperature data, as well as figures on grain supply, export and import, 
and mortality in the 1810s. 

Air Temperatures 

In view of the fact that monthly means of temperature tend to fluctuate about ‘normals’ and 
‘normals’ based on a Series of years fluctuate themselves, both in magnitude and sign, I propose 
to look at temperature differences between the two years flanking 1815, and compare these with 
the parallel differences between 1884 and 1882. As too little is known about the exposure of 
thermometers in the early nineteenth century and the effects of the particular exposure, I assume 
that treating differences of adjacent years will minimize the effects of exposure differences from 
present-day standards. 

Since Tambora ejected a much greater mass of material than Krakatau in 1883 (Appendix 2, 
No. 2; Sigurdsson and Carey, this volume) and, since Tambora more severely effected parts of 
Europe and North America (Post 1977), I expect that the air-temperature reduction in 1816 
should have been greater than in 1884 - unless other meteorological processes overshadowed the 
effects of volcanic particles, if any. 

Air-temperature measurements in the Baltic Region covering both the 1810s and 1880s are 
available for Copenhagen, Gothenburg, Stockholm and Uppsala, but I will also consider data for 
Trondheim. Baltic stations that have data for the 1810s only are St. Petersburg and Vé6yri 
(63°08’N, 22°14’E, near Finland’s west coast). Temperatures for April through September for 
Stockholm and St. Petersburg are shown in Table 1; those of other stations will be found in 
Appendix 1. Although the year 1816 was called a ‘year without a summer’, data on the spring 
months are included because of their importance for the growth of young plants. The tables cover 
several years of the 1810s, on the one hand, and the years 1882 and 1884, on the other - where 
available. Basel’s data are included for comparison’s sake. 


Table 1: Mean Monthly Temperatures (°C), April through September, of Stockholm, St. Petersburg 
and Basel (the Latter for Comparison), and Differences of 1816’s Temperatures from 
Those of Some Previous Years as Well as Differences of 1884 from 1882.! 

Mean of 
Year(s) A M J J A S year 
1809 -0.2 11.3 14.4 17.6 18.3 13.0 4.8 
1812 -1.2 6.5 13.9 14.5 16.6 | 4.1 
1814 a3 7/5 12%2 19.2 16.7 10.6 4.4 
1816 2, 6.0 iss 18.5 14.7 11.8 4.9 
1817 2S 10.7 14.2 17.3 15:2 12.6 5.7 
1882 3.9 9.8 14.3 17.0 17.4 1322 6.7 
1884 Del 8.1 12.6 16.8 14.6 14.1 339 
(1816)-(1814) -2.6 -1.1 +2.9 -0.7 -2.0 +1.2 +0.5 
(1884)-(1882) =e -1.7 -1.7 -0.2 -2.8 +0.9 -1.2 
(1816)-(1809) +2.9 -5.3 +0.7 +0.9 -3.6 -1.2 +0.1 
(1816)-(1812) +3.9 -0.5 +1.2 +4.0 -1.9 +2.7 +0.8 
St. Petersburg 
1809 Sew) 7.4 14.8 16.4 16.0 jl es) 12 
1812 1e3 6.8 13:7 17.9 19.6 4 2 
1814 1.6 6.0 15.3 20.7 17.0 10.3 2.6 
1816 3:6 75 15:1 19.1 14.7 12.4 35 
1817 2.0 11.6 14.0 19.1 WES) 8.1 5 i 
(1816)-(1814) +2.0 +1.5 -0.2 -1.6 -2.3 -2.1 +0.9 
(1816)-(1809) +5.3 +0.1 +0.3 +2.7 -1.3 +1.1 +203 
(1816)-(1812) +2:.3 +0.7 +1.4 +1.2 -4.9 +4.7 +1.0 
(1816-1814) -2.4 +0.1 -1.4 -3.4 -1.4 +0.9 -0.7 
(1884)-1882) -1.7 +0.6 -1.7 +3.1 +2.4 +1.6 +0.3 

' Stockholm - Hamberg (1906, pp. 13-14). Also Historisk Statistik for Sverige, (1959, Volume II, pp. 3-4). 
St. Petersburg - Wahlén (1881, p. 16). 
Basel - Bider et al. (1959, pp. 407-409). 
A = April, M = May, etc. 

Differences for the means for the April-September seasons for all stations are shown in Table 2. 
All the differences (1816)-(1814) are negative. So are the differences (1884)-(1882), except for 
Basel. In Trondheim’s case the “‘Tambora difference’ is more negative than the ‘Krakatau 
difference’; in the case of Stockholm and Uppsala, the Krakatau difference is more negative; in 
Gothenburg’s case, the two are about equal. Hence, the aforementioned sets of differences do not 
indicate a marked Tambora effect at all the stations. Moreover, as can be seen (Table 2), at all 
Baltic stations the spring-summer season of 1812 was colder than that of 1816. Stothers (1984, 


Figure 4) shows, too, that the stations situated in northwestern and northern Europe along the 
5°C annual isotherm were colder in 1808, 1809 and 1812 than 1816. Evidently the years 1807-11 
were not characterized by sizable volcanic eruptions: Lamb’s (1972, Table 10.3) Dust Veil Index 
for the aforementioned four years is much lower than for Tambora, so there is little reason to 
attribute the coldness of 1809 and 1812 to volcanic particles. Because the first two or three 
decades of the nineteenth century produced several cold years or winters, it is doubtful that the 
minor coldness of 1816 relative to 1814 in the Baltic Region was a sequel of Tambora. 

Table 2: Mean Seasonal (April through September) Differences of Air Temperatures (°C). 

Station (1816) (1884) (1816) (1816) 

(1814) (1882) (1809) (1812) 
Trondheim -0.73 -0.48 -0.63 +0.68 
Gothenburg -0.37 -0.38 -1.08 On 5 
Copenhagen -0.42 -0.20 -1.17 +0.30 
Uppsala -1.0 (?) ='-34() ras a 
Stockholm -0.38 =i18 i172 -0.93 +1.57 
Voyn -0.30 seed small + 1.92 
St. Petersburg -0.45 Ese tele +0.9 
Basel -1.27 Ol72. 

In consequence of the appreciable fluctuations of temperature from one year to the next, also in 
periods when no important volcanic eruptions occur, I consider comparison of temperatures does 
not yield convincing results regarding possible volcanic effects (Appendix 2, No. 3) - at least not 
in the period covered here. Harvest data can supply more trustworthy inferences. 

Grain Supply, Mortality - Some General Remarks 

The only country of the Baltic Region for which both harvest estimates and grain import-export 
data are to be had for the 1810s is Sweden. Finland comes next, since data are available on the 

difference between exports and imports as well as ‘qualitative’ estimates of the degree of success 
of the harvests. 

In pre-industrial societies of Europe death rates rose to a maximum either in the year of a serious 
crop failure or in the year following; in other cases, the rise was, or may have been, indirectly 
due to a poor diet, (e.g., the consumption of tree bark, roots, etc.) which promoted endemic 
diseases. The consequences of a crop failure were especially grave when a dearth-stricken country 
could not obtain supplies from abroad (e.g., in the case of wars like that of Norway in 1812-13: 
see below). 

Sweden - Grain Supply 

Figures on grain harvests in Sweden in the 1810s are published in Historisk Statistik for Sverige 
(1959, Volume II, p. 46, Table E16). Data on import and export of grain are listed in Tables 1 
and 2 of the monograph Spannmdlshandel och Spannmdlspolitik i Sverige 1719-1830 (Grain Trade 


and Grain Policy in Sweden 1719-1830) by the Swedish economist K. Amark (1915, pp. 354-355 
and 356-357). The import-export quantities of flour are so small that they were neglected here. 

Export-import data were collected through the Customs Authority and appear to be fairly reliable. 
Because the compilers-editors of the Historisk Statistik (Volume II, 1959, p. 17°) warn that the 
harvest estimates are not very reliable, I base my comments on the import-export data. These data 
are independent of the harvest-estimate figures. All the aforementioned data are listed in Table 3. 

Table 3: Estimates of Grain Harvest and Data on Excess of Import Over Export of Grain for Sweden in 
the 1810s in Thousands of Metric Tons.’ ? 

A B Cc D 

1811 512 24 536 
1812 473 42 515 
1813 530 96 626 
1814 568 62 630 
1815 581 21 602 
1816 527 22 549 
1817 545 19 564 
1818 473 54 S27 
1819 532 40 a12 
1820 654 1 655 

' A - year, B - grain harvest, C - excess of import over export of grain, D - B+C. 

2 Grain harvest Sources: Historisk Statistik for Sverige (1959, Volume II, Table E16, p. 46); 
Import-export: Amark (1915, Tables 1 and 2, pp. 354-355 and 356-357). 

The excess of imports Over exports was very high in 1813 (96,000 metric tons). In fact, a 
reference to Amark’s Table 1 indicates that not since 1785 (winter 1783-84 was rather cold, a 
fact mentioned in Appendix 2, No. 3) were imports as high as in 1813. Amark (1915, pp. 
106-108) repeatedly states that 1812 was a year of crop failure (“missvaxt’) (Appendix 2, No. 4). 
This failure necessitated increased imports the following year. No mention is made of a dearth 
in 1816 and 1817. Table 3 shows, indeed, that in the two post-Tambora years, imports were but 
one-fifth to one-quarter of the imports of 1813. 

As has been stated, the spring and summer of 1812 were cold in the Baltic countries. At 
Stockholm the mean temperature of April, May and July was 6.6°C, whereas the corresponding 
mean of 1816 was 9.1° and that of 1817, 10.2°. Thus the growing seasons of the two post- 
Tambora years were warm. Similarly, Gothenburg’s temperature data (Historisk Statistik, 1959, 
Volume II, p. 6) indicate that in 1816 and 1817 spring and summer were warmer than those of 
1812, though the differences were smaller than in Stockholm’s case. 


The economic historian Sommarin (1917) published a monograph The Economic Development of 
the Scanian Agriculture (1917; in Swedish). As the title indicates, the emphasis of the work is 
on Scania, Sweden’s southern region. Sommarin classifies the harvests of Malméhus and 
Kristianstad, the two southernmost counties, into seven groups: 0 to | = crop failure, ..., 4 = 
average, ..., 6 = abundant, rich. In a Table (pp. 208-209) he puts the harvests of both 1816 and 
1817 into group "6", that is, abundant. In contrast, the harvest of 1811 is estimated by him as 
"3.4" and that of 1812 as "5.1"; the three years 1813-15 were considered "a little above 
average". I will discuss 1811 later under Denmark. 

A commentary on the meteorological background of the below-average harvest in southern 
Sweden in 1811 is offered by rainfall data for Lund, the university city in Malmoéhus. 
Precipitation records at Lund began in 1748 at the Astronomical Observatory of the University. 
The rainguage was placed on the Observatory’s roof, 20 m above ground, during the period 
1775-1867 (Tidblom 1875-76, p. 65). This is an objectionable exposure, but the results of the 
measurements are still of value in comparing one year with another. Cumulative rainfall for April 
through July (the monthly figures are found in Tidblom 1875-76, pp. 66-68), the main growing 
season, follow: 

Cumulative Rainfall (mm) 

Year at Lund, April through July 
1811 98 
1816 247 
1817 158 

Clearly, the rainfalls of 1816 and 1817 were much more abundant than that of 1811. 

Sweden - Mortality 

Volume I of the Historisk Statistik for Sverige 1720-1967 (1969) includes population data for 
Sweden. On p. 93 the death rates for 1811-20 are given (see Table 4, with similar data for the 
other Baltic countries, excepting Russia). 

In the 1810s, 1816 had the lowest mortality , whereas 1812 had the highest rate (Table 4). One 
of the reasons for the low mortality in 1816 is that the largest population-increase rate of the 
decade occurred from 1815 to 1816 (incidentally, 1815 was a good crop year). In any case, 
mortality was not high in the two post-Tambora years. That the crop failure of 1812 did not 
cause famine or high mortality was largely due to the fact that Sweden was able to import large 
quantities of grain (partly from Swedish Pomerania and from Danzig), being a kind of trading 
centre for Polish grain that was conveniently shipped on the Vistula. Swedish Pomerania was 
adjacent to the southern coast of the Baltic. Sweden’s ability to import large quantities of grain 
in 1813 was mainly due to her siding with the British-Russian coalition against Napoleon. 

Norway - Grain Supply 

No crop-production or grain import-export data are available for Norway for the 1810s. 
However, Norway suffered a grave crop failure in 1812 (e.g., Steen 1933, p. 339; Drake 1969, 
p. 60), and, unlike in Sweden, 1813 was a year of famine. The British blockade of Norway 


Table 4: Mortality in the Baltic Region, 1811-20, in %o of Population. ' 

Year Denmark Finland Norway Sweden 
1811 24.4 30.8 25 28.8 
1812 27.0 23.9 ZA 30.3 
1813 22.8 273 29.5 27.4 
1814 24.7 26.6 22.6 25a 
1815 21.6 26.0 29.8 23.6 
1816 20.7 23.4 19.4 22 
1817 19.0 24.1 I W/ECE 24.3 
1818 18.9 251 19.1 24.4 
1819 19-5 21:3 19.7 27.4 
1820 20.9 2523 18.9 24.5 

' Sources: Denmark: Andersen (1973, p. 300); Finland: Turpeinen (1979, p. 108); 
Norway: Drake (1969, p. 194); Sweden: Historisk Statistik for Sverige 
(1969, p. 93). 

virtually cut off the country from foreign sources, including Denmark which used to supply about 
25% of Norway’s grain needs (Derry 1965, p. 488). Neither 1816 nor 1817 were years of dearth. 

The historian Steen (1933, p. 339) writes that 1812 was the blackest year that the poor of Norway 
had to experience. Spring was late. In many places, the grain could not be sown before June. 
Another historian, Mykland (1978, p. 239) states that on 9-10 August there was a strong 
nightfrost, in consequence of which the crop was totally destroyed in the northern part of the 
country. Returning to Steen’s account, in some parts of Norway in September snow covered the 
unripe grain. 

The famine was so severe that Prince Frederick of Hessen, Norway’s Vice-Governor, wrote King 
Frederick VI, in Copenhagen recommending that the British be asked for an armistice (Mykland 
1978, p. 228) "not to have to succumb to that most frightful scourge, famine". The King did not 
accept the recommendation. From May to autumn 1813, bread riots broke out in many areas of 
Norway (Lindval 1952, p. 129). 

A reference to the temperature data of Trondheim (Appendix 1) shows how much lower the 
temperatures were in 1812 compared with those of 1811 and with some of 1816. 

Norway - Mortality 

Drake (1979, p. 291) publishes a diagram showing the mortality rates in dioceses of Norway and 
in the country as a whole. The diagram shows that mortality reached a peak in 1813 in the 
Akershus and Trondheim dioceses (Akershus includes Oslo), with another peak in 1809 - also a 
cold year. The fact that the curve for the country exhibits a peak in the same years as Akershus 
and Trondheim dioceses is due to the fact that those dioceses comprised 67% of Norway’s 
population (Drake 1979, p. 290). As to the peak for Akershus, Drake (1969, p. 71) states: "A 
number of major peaks in the death rate [in Norway] ... occurred in years when the grain harvest 


failed over large parts of the country and it is noticeable in each that the Akershus diocese, where 
the diet was based more on grain than elsewhere, suffered particularly badly." 

Drake’s diagram shows that there was no rise in mortality in either 1816 or 1817. A reference 
to Trondheim’s temperature data (Appendix 1) indicates that April, May and July 1816 and 1817 
were warm relative to 1812 - just as at Stockholm. 

Denmark - Grain Supply 

Denmark was a grain-producing and exporting country during this period. According to 
information received from Professor Claus Bjorn, historian of agriculture at the University of 
Copenhagen, no harvest figures and no complete sets of data are available on exports to various 
countries in the 1810s. The only data available relate to the export of grain to England from 1816 
on. The Danish economist Falbe-Hansen (1889, p. 16) publishes the following grain-export 
figures for 1816-25: 

Year Quarters 
1816 14,900 
1817 149,000 
1818 342,000 
1819 123,000 
1820 147,000 

These quantities are small, but indicate that there was no shortfall of harvest either in 1816 or 
in 1817: the export of 1817 was made possible by the harvest of 1816. 

Professor Bjern has drawn my attention to an interesting volume of 1811-38 records of a 
tenant-farmer Seren Pedersen. Pedersen lived at Havrebjerg in southerwestern Sjzlland. He was 
not only a farmer but also the Executive Officer of his parish. As an Executive Officer, he kept 
accounts and records of events, at the end of each year recording his thoughts on the character 
of the year’s weather, crops, prices, etc. In his reflections on 1816 (Pederson 1983, pp. 214-215) 
he writes: 

"This year was again a lovely fertile year, not only on account of the quantity of 

fodder but also because the kernels were big. This was important for our country 

as this commodity was very much coveted in foreign lands. The price was high so 

that the monetary situation of the country has improved and one assumed that the 

bad times of war [the Napoleonic Wars] are over. The rate of exchange was nearly 

par and one thought that the good old times have returned and this not only because 

of our fortunate grain trade with other countries where there has been a general 

crop failure. Denmark has been selling this year for many millions of rigsdaler [old 

Danish monetary unit] grain to other nations." 

Probably his reference to crop failure in other countries and the demand for Danish grain reflect 
"the year without a summer" in western and central Europe. 


In his contemplations on 1817 (Pedersen 1983, pp. 223-224) we read the following: "This year 
was, I suppose, not as happy as the previous year, neither with respect to fertility nor the prices 
obtained for our produce. But, it was not a year of famine, not a barren year, rather an average 

In contrast, in his review of 1811, he states (Pedersen 1983, pp. 168-170): "The year 1811 ... 
a hard and barren year. There was a shortage of both grain and hay because of the too dry 
summer. This led to very high prices, even of commodities of poor quality ... The public 
authorities conducted enquiries how much each peasant had reserves so that they can distribute 
the food, if necessary." 

We see from Pedersen’s reflections that 1816 and 1817 were good to average crop years, whereas 
1811 was a year of drought (Appendix 2, No. 4) and crop failure, at least in a part of Denmark. 
A reference to Copenhagen’s air-temperature data in Appendix 1 shows that 1811 was, indeed, 
a warm year. As to rainfall, this was measured in Copenhagen in the 1810s on the top of a 36-m 
high tower (The Round Tower) in the city centre. However, the exposure of the gauge was 
unsatisfactory. Some reliance can be placed in the precipitation measurements at Lund, southern 
Sweden, which, as has been pointed out previously is but a short distance from Sjzlland across 
the Danish Sound. We have mentioned also the Lund rainfall data for spring-summer 1811 
showing that rainfall was low - as stated by Pedersen for southwestern Sjeland. 

That Pedersen’s statement concerning a poor harvest in 1811 applies not only to his 
neighbourhood, also follows from a statement in minutes of a meeting of "Sjelland’s Provisioning 
Committee" on 31 August 1811. Dr. Helle Linde (personal communication 1988) of the 
Copenhagen City Archives has kindly brought to my attention a passage in the minutes which 
translates as follows: "It is known that the [price of] grain fell substantially during the first 
summer months when the crops looked promising. But it has turned out soon that the harvest of 
this year will not be better and that it will be even less satisfactory than the harvest of the 

The listed prices of grain and other commodities in the Danish newspaper Danske Statstidende 
for 1809 and 1811 indicate a clear rise compared with spring 1811 and those of 1809 
(Appendix 2, No. 5). 

Denmark - Mortality 

Figures for Denmark (Andersen 1973, p. 300), and parallel data for Finland, Norway and 
Sweden show that mortality in Denmark was relatively low in both 1816 and 1817, whereas it 
was high in 1812 (Table 4). 

Finland - Grain Supply 

Data are available on the excess of export over import of grain as from 1812, collected by the 
Customs Authority. Additionally, ‘qualitative’ harvest estimates are recorded in reports of the 
Provincial Governors to the ‘Senate’ (the semi-autonomous Government of Finland under Tsarist 

Data on excess of exports over imports have been published by the Finnish agricultural historian 
Soininen (1974, Table 24, pp. 190-191), and the Provincial Governors’ reports have been studied 
and summarized by Johanson (1924, see especially pp. 82-85). The export-import figures for 
1812-20 are listed (Table 5a,b). The large excesses of imports in 1813 and 1819 suggest poor 


harvests in 1812 and in 1818. These inferences are in close agreement with estimates of the 
Provincial Governors who state that 1812 was a year of severe crop failure, whereas 1818 was 
a year of crop failure. 1815 was likely to be a good year and 1816 and 1817 did not excel, but 
the Governors do not classify them as years of crop failure. Table 5b gives the amounts of grain 
exports to Sweden from Finland, showing that during 1816-18 exports were appreciable, whereas 
in 1819 (after the crop failure of 1818) exports to Sweden were reduced. 

According to Johanson (1924, pp. 82-83), the spring of 1812 was very cold - as were all the 
Baltic countries then (see section on Air Temperatures above, and temperature data in 
Appendix 1). He adds that the summer of 1812 was cool and rainy, leading to a severe shortfall 
of the harvest. The same author states that the harvest of 1816 was normal, except in the Viipuri 
(Viborg) and Vaasa (Vasa) counties. In 1817 the spring and summer were rainy, but the crops 
did not suffer. 

Finland - Mortality 

In Table 4 are listed Finnish mortality data as published by Turpeinen (1979, Table 1, p. 108). 
Mortality was relatively high in 1811, and, after a major drop in 1812, it rose in 1813. Another 
rise, compared with the flanking years took place in 1819. Presumably, the peaks are related to 
crop failures of the previous years. On the other hand, we note that mortality was low in 1816, 
though a slight rise is seen in 1817 and 1818. 

Table 5a: Excess of Exports over Imports of Bread Cereals and Oats, in Thousands of Hectolitres, 
from Finland, and Classification of Grain Harvests by Provincial Governors.' 

Year Bread cereals Oats Total Harvest 
1811 aes a: = crop failure 
1812 49.5 10.9 60.4 severe crop failure 
1813 -179.6 10.5 -169.1 oe 

1814 -7.3 -5.0 -12.3 abundant crop 
1815 13:2 -6.5 6.7 eas, 

1816 98.5 0.1 98.6 oe 

1817 -66.5 Si /iP -73.7 = en 

1818 -141.5 -12.5 -154.0 crop failure 
1819 -274.7 -31.4 -306.1 nate 

1820 -57.5 -9.1 -66.6 

' Sources: Export-Import - Soininen (1974, Table 24, pp. 190-191). Harvest estimates - Johanson 
(1924, pp. 82-85). 


No figures on grain production, nor on grain export-import are available, but there is information 
on years of famine and, in some cases, on years when the crop failed. References to famines in 
Russia can be found in a bibliography prepared by D.R. and V. Kazmer (1977) titled Russian 
Economic History, a Guide to Information Sources. The most comprehensive English paper on 
famines in Russia is that by Kahan (1968). Other articles in English are by Dando (1981) and a 
brief review paper by Robbins (1979) in The Modern Encyclopedia of Russian and Soviet History. 
All three list dates and brief remarks on the degree of severity of the famines. 


Table 5b: Export of Bread Cereals, Oats and Peas 
from Finland to Sweden Alone, in Metric 


Year Export 
1815 16,300 
1816 73,000 
1817 53,000 
1818 62,000 
1819 16,000 
1820 13,000 

' Source: Johanson (1924, p. 165). 

The most detailed list (Kahan 1968, pp. 367-375) includes a good discussion. As to the early 
nineteenth century, Kahan says that grain yield was low in 1812 near Moscow and in Siberia, and 
that in 1817 there was a drought in the non-black soil region, causing a rise in prices. Compared 
with the specifications given of other dearth years, 1817 must have been a light case. In any case, 
no dearth or famine is reported for 1816. Neither Dando’s nor Robbins’ lists mention dearth in 
1817. Sorokin (1975, p. 179) writes that in 1812, 2'% million rubles were appropriated by the 
Government for the purchase and distribution of bread to the starving population of Moscow 
province, and that in 1813 the provinces of Kaluga and Smolensk were granted 6 million rubles 
for similar purposes. Such shortages were, in all probability, a sequel of the cold year 1812. No 
mention is made of dearth in 1816 (Appendix 2, No. 6). 

In addition to the above works in English, several Russian encyclopedias published in Tsarist 
times toward the end of the nineteenth century furnish rather long articles (unlike the Great Soviet 
Encyclopedia) on famines in Russia and Europe. None mentions famine in Russia in 1816 and 
1817 though one remarks on famine in Germany in 1817. 

Marshall (1833 p. 95) quotes figures on grain exports in 1791-1825 from northwestern and 
northern Russian ports (Libau = Liepaja, Riga, St. Petersburg and Archangel). While quantities 
were not large, there was a manifold increase from 1816 to 1817. As it was the custom of 
Governments to prohibit exports in years of dearth, there could not have been serious shortages 
in the Empire. On the other hand, the column for Riga shows that in 1818 and 1819 no exports 
were allowed. 

Grain Prices 

Prices are not always good indicators of harvests. Governments, including local ones, used to 
control prices, subsidize sales to the poor and, in some cases released stocks earmarked for 
military purposes to the general public. For example, in Norway in 1813, quantities of grain were 
taken by the Regent from army stocks. Nevertheless, I shall make an exception and quote Abel 
(1974, pp. 318-319) who points out that from 1815 to 1817 the rise in the price of rye in Sweden 
(rye being a staple food of the population) amounted to only 20-30%. At Danzig, the principal 


grain-trade centre of the Baltic Region, which was selling to countries near and far, the rise was 
18%. These figures are to be compared with the following figures for price rise in some of the 
Tambora-stricken western European countries (e.g., England 150%, France 185%). 

Why the Baltic Region Was Not Reached by Large Numbers of Volcanic Particles from 
Tambora’s Eruption 

Why was the Baltic Region not affected by Tambora’s eruption of 1815? It is puzzling, indeed, 
that while western and central Europe south of 50-55° were gravely hit, the Baltic Region and 
European Russia, including its south, were not. Because the acidity accretion of 3.8+0.1 p equiv. 
H* per kg ice in the 1816 layer of central Greenland (71°) really is relatively small, perhaps 
aerosol plumes of Tambora crossing the Baltic Region were too ‘thin’ to produce a noticeable 
impact on air temperatures and agriculture. 

The following comments are based in part on computed data and in part on speculation. 
Speculation is inevitable for 1816 since we have no other meteorological data than surface air 
temperatures at an admittedly small number of stations. 

Direct Solar Radiation (DSR) 

Table 6 lists the amounts of DSR reaching a horizontal surface at the top of the atmosphere in 
the summer and in the winter half-years respectively (summer half-year: 21 March to 
23 September). The figures were computed by List (1958, p. 418) on the assumption that the 
solar constant equals 1.94 cal cm? min’. 

I will compare the figures for Latitude 60° with those for Latitude 45°. Oslo, Stockholm, 
Helsinki and St. Petersburg all are situated near the former. 

In the summer half-year the DSR reaching a horizontal surface at the top of the atmosphere at 
Latitude 60° is about 10% less than Latitude 45°. However, when account is taken of the 
depletion of DSR along the longer atmospheric path to the stratosphere at 60°, the difference 
between the two latitudes is bound to be greater. This means that in the atmosphere of the higher 
latitudes any volcanic particles have less DSR to scatter back to space. If we assume that the 
two latitudes compared have, hypothetically, the same size distributions of particles, and that they 
have the same physical-chemical characteristics of the particles, then the reduction of air 
temperature in the higher latitudes will be less than in the lower mid-latitudes. 

In the winter half-year the difference in DSR at the top of the atmosphere is 47%, and the 
difference at stratospheric levels is much greater. We can expect, consequently, that in the winter 
half year volcanic particles in the higher latitudes will scatter back to space very little DSR and, 
concomitantly, the diminution of air temperature will be very small. Even in the lower mid- 
latitudes, where the year 1816 was ‘without a summer’, the winter shows minor effects. 

Volcanic Particles 

Kondratyev et al. (1983) provide a useful review of our knowledge of the nature and 
characteristics of volcanic particles. In the case of powerful eruptions both particulate and gaseous 
matter are injected into the stratosphere which the atmospheric circulation can carry great 
distances. (We are not concerned here with the ‘close fall-out’ of heavy ash.) The gaseous matter 
is predominantly SO, which undergoes with time (~transport by winds) photo-oxidation 


Table 6: Direct Solar Radiation at the Top of the Atmosphere (in cal 
cm’) in Summer and Winter Half-Years on the Assumption that 
the Solar Constant Equals 1.94 cal cm? min-1. (Summer half-year: 
21 March to 23 September’). 

Summer Winter 
Latitude half-year half-year 
70° 134,540 13,040 
60° 144,610 32,610 
50° 156,030 56,980 
45° 160,790 69,360 
40° 164,620 81,510 
305 169,220 104,570 

" Source: List (1958, p. 418). 

converting it into sulphuric acid. The formation of sulphuric acid involves the heterogeneous 
nucleation on small mineral particles and homogeneous nucleation, the latter if a high degree of 
supersaturation of H,SO, vapour prevails (e.g., Hofmann and Rosen, 1983b, p. 327). It is 
probably correct to say that high supersaturation conditions can occur in the stratosphere rather 
than in the troposphere. In the conversion process mineral particles coated with sulphuric acid 
and sulphuric acid droplets are produced (Appendix 2, No. 7). In addition to the process of 
formation itself, the fact is of special importance to our considerations that sulphuric acid 
particles/droplets tend to grow with time. As to the sulphuric-acid coating growth, I have quoted 
(Appendix 2, No. 1) Mossop’s communication concerning the volcanic particles from Agung’s 
eruption; see also the papers of Hofmann and Rosen (1982, 1983a and b) concerning particles 
from other eruptions. 

As the coating grows and the particles become larger, their terminal velocity of fall increases. 
However, the process of enhancement of terminal velocity is counteracted by the increasing air 
density. The terminal velocity is given by Stokes’ equation: 

2a rp 
Ve ( - ! ; (1) 
&.¥ p 
provided that the Reynolds number, Re, 
2QV 45 GE) ip 
= | -1 | (2) 
y 2s Nap 

is appreciably less than 1, a condition that is amply satisfied by our particles. In the above 
equations a is the radius of the supposedly spherical particles or droplets and p is the density of 
the particles/droplets. The other symbols are standard. 

In the troposphere the terminal velocity of the ‘large’ particles that have fallen out from the 

stratosphere will decrease (or, continue decreasing) as they carry on falling, unless the growth 
of the particles outweighs the effect of increasing air density. However, the chances of growth 


of the sulphuric-acid coating, or that of the sulphuric acid droplets, are less favourable in the 
troposphere than in the stratosphere since most of the SO, of large eruptions is injected into the 
stratosphere. Assuming, for purposes of illustration, no growth of coating of the particles in the 
troposphere, and no coagulation and or washout by rain, a particle of radius of 1 wm and a 
density p = 2 g cm” will have in the International Standard Atmosphere at 10 km (T = -50°C, 
p= 0413 x 107 ecm, vy ="0:092 ems) V = 0.117 cms", while at 5 km(T = -17.5°C, p 
= 0.736 x 10°g cm’, vy = 0.105 cm? s") V = 0.057, that is, half the terminal velocity at 10 km. 
Thus, in a climatic zone where there is little precipitation (and, if other assumptions hold), the 
concentration of volcanic particles will tend to increase with time in the troposphere - at least for 
a while. 

On their way to Europe (and North America) the Tambora particles had to cross the subtropical 
high-pressure belt of the northern hemisphere. In many areas of that belt there is very little 
precipitation, especially in summer. Once moved into the middle latitudes, the tropospheric air 
that became enriched by volcanic particles in the high-pressure belt, is liable to lose many of the 
particles by washout, a process that is orders-of-magnitude more efficient than the process of ‘dry 
fall-out’. Additionally, the stratospheric air reaching the mid- and higher latitudes from the 
high-pressure belt was likely to have become depopulated of small particles as these grew and 
fell into the troposphere. 

If our speculations are essentially correct, then the Baltic Region should have been crossed by 
air poor in sulphuric-acid particles and these relatively few particles could not scatter back to 
space more than a very small fraction of the DSR. 

Abel (1974, pp. 318-319) points out that Tambora effects decreased from south to north and from 
west to east in Europe. How did it happen that southern European Russia was spared? After all, 
the latitudes of southern Russia are roughly the same as those of western and central Europe that 
bore the brunt of Tambora’s eruption. I suggest that a consideration of the direction of 
tropospheric winds resolves the enigma, assuming that the directions were much the same as in 
recent decades. The tables in the Handbook of Geophysics and Space Environments (1965, 
pp. 4-48) show that in summer, in the Longitudes of western and central Europe in Latitudes 40° 
and 50°, in the higher troposphere, the winds flow from west to west-southwest. These winds 
would ‘carry’ to southern Russia air largely depleted of volcanic particles by precipitation in the 
more western areas of the continent. Precipitation data for a few stations in western Europe show 
that the summer of 1816 was definitely not dry (Table 7). 

The fact that the low mid-latitudes of western and central Europe suffered greatly but not the 
Baltic Region, despite the relatively short distance between them (from 1,000 to 1,500 km and 
under; in comparison Europe is about 10,000 km from Tambora), suggests that a drastic factor 
cleaned the air of the low mid-latitudes of many of its volcanic particles. Washout by rain is such 
a factor. 

Absorption of DSR by Sulphuric-Acid Particles 

It is worth looking at the possibility that the apparent failure of supposedly numerous volcanic 
particles to reduce temperatures in the Baltic Region was due to absorption of the DSR by the 
sulphuric-acid particles, and that this hypothetical absorption left little DSR to scatter back to 
space. In this context my (Neumann 1973, pp. 96-97) laboratory measurements of the absorption 
spectra of water solutions of sulphuric acid at three high concentrations, viz. 49, 73 and 98% 
(Appendix 2, No. 8), as a function of wave length of radiation are worth considering. The 


Table 7: Precipitation (mm) in 1815, 1816 and 1817 at Some Stations in the Lower Mid-Latitudes of 

Europe. ' 
A M J J A S 


1815 716 113 85 181 116 10 
1816 84 716 109 74 719 54 
1817 6. il 48 105 84 54 

1815 30 29 719 32 15 32 
1816 13 38 54 97 51 63 
1817 1 65 102 59 50 62 

1815 68 58 48 45 45 30 
1816 53 55 60 108 63 55 
1817 3 115 35 108 68 23 

' Sources: Milan, Brunt (1925, p. 277); Paris, Garnier (1974, p. 51); Kew (London), Wales-Smith (1971, 
p. 359). A = April, M = May, etc. 

results, obtained by advanced instrumentation, show that sulphuric acid solutions have very low 
(nearly zero) absorption for light at wave lengths between 0.3 and 1.5 um. Since about 90% of 
the DSR resides in that wave-length range, clearly there could have been no important absorption 
by sulphuric acid particles of DSR. 

The result that concentrated water solutions of sulphuric acid have virtually no absorption in the 
range of DSR, is confirmed by an independent statement of Deirmendjian (1973, p. 293, item 
(iii)). Referring to the physical and chemical characteristics of volcanic particles, he writes: " ... 
the particles may have been composed of nonabsorbing (or very weakly absorbing) dielectric 


I suggest that data quoted in this paper, especially those on grain supply and mortality, indicate 
conclusively that Tambora’s eruption of 1815 had slight effect in the Baltic Region. Additional 
support for this conclusion is rendered by the fact that Post’s monograph The Last Great 
Subsistence Crisis in the Western World (1977) does not quote data indicative of any significant 
effect on the region (Appendix 2, No. 9). 

A plausible explanation for the observation that the Baltic Region showed no perceptible effects 
of Tambora’s eruption is that the air reaching the region was cleansed of volcanic particles by 
precipitation in western and central Europe. I think that much the same principle applies to 
southern European Russia which, like the Baltic Region, was not harmed by the eruption. 



I thank the following scientists for their ready assistance: Professor Claus Bjgrn, Institute of 
History, University of Copenhagen, for information on literature on agriculture in Denmark in 
the 1810s; Drs. H.B. Clausen, D.A. Fisher and C.U. Hammer, Department of Glaciology, 
Institute of Geophysics, University of Copenhagen, for assistance with literature and comments; 
Dr. C.R. Harington, Coordinator "Climate in 1816" Meeting, Paleobiology Division, Canadian 
Museum of Nature, Ottawa, Canada, for his invitation to prepare this paper; Lars Landberg, 
graduate student, Department of Meteorology, Institute of Geophysics, University of Copenhagen, 
for translation from the Danish; Dr. Helle Linde, Copenhagen’s City Archive, for information 
and some passages from documents relating to 1811; Dr. Jarl Lindgrén, Population Research 
Institute, Helsinki, for assistance with literature; Susanne Lindgrén, Adviser, Secretariat, Nordic 
Council of Ministers, Copenhagen, for past and present discussions on the subject of 1816 in the 
Baltic Region and for literature references; Torben Pedersen, Department of Meteorology, 
Institute of Geophysics, University of Copenhagen, for translation from the Danish; Dr. Arvo 
M. Soininen, formerly Lecturer in Agricultural History, University of Helsinki, for comments 
and for copies from the literature; Dr. Oiva Turpeinen, historical demographer, Institute of 
History, University of Helsinki, for comments and literature; Dr. Cynthia Wilson, Gillingham, 
Kent, England, for her readiness to present the substance of this paper at the "Climate in 1816" 
Meeting, Ottawa, Canada, June 1988. 

The following libraries (and their librarians) were helpful with literature: Copenhagen: The 
Danish Veterinary and Agricultural Library; The Royal Veterinary and Agricultural University; 
Library of Denmark’s Statistics; Royal Library; The University of Copenhagen’s (a) Library of 
the Institute of Economic Research, (b) Library of Humanities, Fiolstrede and Nijalsgade 
divisions, (c) Library of Slavic Philology; G6ttingen: University of Gottingen’s Library; 
Helsinki: University of Helsinki’s Library/Slavica. 


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Bider, M., M. Schiiepp and H. von Rudloff. 1959. Die Reduktion der 200 jahrigen Basler 
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Brunt, D. 1925. Periodicities in European weather. Philosophical Transactions of the Royal 
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Deirmendjian, D. 1973. On volcanic and other particulate turbidity anomalies. In: Advances in 
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Dettwiller, J. 1981. Les températures annuelles a Paris durant les 300 années. La Météor 
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Hammer, C.U., H.B. Clausen and W. Dansgaard. 1980. Greenland ice sheet evidence of 
post-glacial volcanism and its climatic impact. Nature 288:230-235. 

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Appendix 1 

Air-temperature data (°C) for Copenhagen, Gothbenburg, Kiev, Trondheim, Uppsala, Véyri and 
Warsaw for some of the years of the 1810s and for 1882 and 1884 - where available. Although 
Kiev and Warsaw are not part of the Baltic Region, their data are cited to reinforce the 
observation that the effects of the Tambora eruption seem to have diminished northward and 
eastward in Europe. This observation, made previously by Abel (1974, p. 319), appears to be 
supported by the temperature data. For example at Kiev, 1814 was about as warm as 1812 
relative to 1816; at Warsaw 1816 was substantially warmer than 1812. V6yri (western Finland) 
was considerably colder in 1812 than in 1816. Hence, it is uncertain that any coldness of 1816 
relative to 1814 was due to Tambora. In the tables, A = April, M = May, etc. 

Mean of 

Year(s) A M J J A S year 

1811 4.8 13.6 lw fee) 19.0 17.0 13.8 8.8 
1812 2.4 9.4 14.8 14.7 16.6 12.1 6.6 
1814 6.0 7.8 1S, 17.8 16.2 12.8 6.4 
1816 5.0 8.1 iS ey 17.0 15.0 1320 6.8 
1817 4.3 11.0 14.8 Se 7/ 15.7 14.9 7.9 
(1816)-(1814) -1.0 +0.3 0.0 -0.8 1-2 +022 +0.4 
(1816)-(1811) +0.2 -5.5 -3.8 -2.0 -2.0 -0.8 -2.0 
(1816)-(1812) +2.6 -1.3 -1.1 toed -1.6 +0.9 +0.2 
1882 6.9 11.0 14.6 17S 16.2 14.3 8.4 
1884 4.8 10.9 14.4 17.5 16.7 14.9 8.2 
(1816)-(1814) -1.0 +0.3 0.0 -0.8 =A, +0.2 +0.4 
(1884)-(1882) -2.1 -0.1 -0.2 0.0 +055 +0.6 -0.2 
(1816)-(1811) +0.2 -5.5 -3.8 -2.0 -2.0 -0.8 -2.0 
(1816)-(1812) +26 -1.3 = +2.3 -1.6 +0.9 +0.2 


Year(s) A M q 4 A S year 

1811 4.6 its ie 16.6 19.8 fei | 13.2 8.2 
1812 Dal 10.2 14.7 15.0 173 Ld 6.3 
1814 6.2 9.3 14.4 18.2 16.9 12:5 6.0 
1816 33 8.5 14.8 18.4 15.9 12.4 6.5 
1817 4.3 11.1 15.4 16.5 15.6 15:3 Vs 
1882 6.1 11.4 15.3 17.6 16.8 14.0 8.2 
1884 5.6 9.9 14.5 17.8 16.6 14.5 8.2 
(1816)-(1814) -0.9 -0.8 +0.4 +0.2 -1.0 “OU, 0:5 
(1884)-(1882) -0.5 -1.5 -0.8 +0.2 -0.2 +0.5 0.0 
(1816)-(1811) +0.7 -4.6 -1.8 -1.4 -1.2 -0.8 -1.7 
(1816)-(1812) oun -1.7 +0: 1 +3.4 -1.4 1.3, +-0,2 

1812 4.9 1257 18.2 22 19.8 13.6 am 
1814 8.6 11.9 7S) 21.6 19.8 13.4 7.4 
1816 Te. 13.6 19.7 18.5 17.9 1537 Vad 
(1816)-(1814) -1.4 Salley) +2.4 -3.1 -1.9 23 +0.3 
(1816)-(1812) ees +0.9 ale -2.7 -1.9 teal 2 

1811 2.8 10.1 ileal 16.2 13:5 10.1 oS) 
1812 0.1 6.4 1h 11.9 13:8 7.8 SiS) 
1814 =e) 6.1 10.4 15.0 13.4 9.4 3.8 
1816 2.8 6.3 11.4 14.1 1L.9 8.7 3.8 
1817 Di 7.9 11.4 43.2 ES) 10.4 4.6 
1882 3.0 8.6 13:2 15.6 15.1 12.0 3:5 
1884 4.2 7.0 11.4 14.4 1525 12.1 6.0 
(1816)-(1814) -2.5 +0.2 TaeO -0.9 -1.5 -0.7 0.0 
(1884)-(1882) So kare -1.6 -1.8 -1.2 +0.4 +0.1 +0.5 
(1816)-(1812) Dal -0.1 +0.3 SIP) -1.9 +0.9 +0.3 

1812 =e 6.3 13:3 13.6 15.0 8.9.0 
1814 SP) 6.8 P27 18.8 16.1 10.9.4 
1816 3-1@) Se7@) 13:9 17.0 13.8 11.0.4 
1817 32 10.8(?) 135 16.8 14.5 P1823 
1882 32 SE) 13.8 16.7 16.6 11.3.8 
1884 2.2 1 11.8 16.1 13.6 12.9.9 
(1816)-(1814) -2.1(?) =i). * 1.2 -1.8 -2.3 +0.1 0.0 
(1884)-(1882) -0.9 -2.2 -2.0 -0.6 -3.0 teil -0.9 
(1816)-(1812) +4.2(?) -0.6(?) +0.6 +3.4 Syl A? +21 +0.4 

' No data have been published for Kiev for the 1880s; no data are available for all the months of 1812 and, 
consequently, the annual mean for that year cannot be given. 


Mean of 

Year(s) A M J J A S year 

1809 -1.3 8.7 i1S),33 Neg, 17.4 10.1 2.4 
1811 -0.3 8.2 17.6 19.0 15.4 OF 3.4 
1812 -2.3 6.8 12.5 15.6 16.7 7.4 Del 
1814 3)72 6.9 14.6 19.7 16.2 9.4 2.8 
1816 ney 8.9 14.4 19.1 14.2 10.5 372 
1817 0.9 10.5 13.0 18.3 14.8 10.5 3.0 
(1816)-(1814) -2.1 +2.0 -0.2 -0.6 -2.0 +1.1 +0.4 
(1816)-(1809) +2.4 +0.2 -0.9 +1.4 -3.2 +0.4 +0.8 
(1816)-(1811) a4 +0.7 230) +0.1 7) +1.4 202 
(1816)-(1812) +3.4 Od +1.9 43.5 O15 Segal ed 

1812 3)-5) 12.4 eS 18.5 C7 11.3 SyoJ/ 
1814 8.7 9.9 15.4 20.1 WH) 10.6 6.0 
1816 6.4 A272 16.8 16.9 16.3 12.6 6.3 
1817 2.4 13.4 N70) 17.5 18.0 12.6 7.1 
(1816)-(1814) -2.3 2.3: +1.4 -3.2 -1.4 +2.0 +-0:3 
(1816)-(1812) 9.0 0.2 0.7 AN6 aA $1.3 +0.6 


Data for Voyri are available only for the years 1800-24. 
No data for Warsaw for the 1880s. 

Appendix 2 

1. By the term ‘developing’ I mean to the process whereby, e.g., the SO, and SO, gases 

emanating from a volcano ‘catalyze’ to sulphuric acid in the course of their transport and 
diffusion. Mossop (1963, 1964, 1965) sampled the stratospheric particles after Mount Agung’s 
(Bali) 1963 eruption and found that they were mostly of a ‘dual’ structure: a mineral core 
around which accreted a water-soluble coating. Mossop (1964) states that the most likely 
composition of the coating was sulphuric acid. Mossop (personal communication 1973) 
indicated that the process of coating growth was such that a year after the eruption the volume 
of coating was about 10 times that of the core (see Neumann 1974, second footnote to p. 
385). It is reasonable to assume that Tambora’s particles developed similarly. In addition to 
SO*, anions, other anions, such as NO, and Cl were involved. 

Langway et al. (1988, p. 7) estimate that the quantity of stratospheric H.SO, from Tambora 

was about an order-of-magnitude greater than that from Krakatau’s eruption. Kondratyev et 
al. (1983) estimate that the stratospheric H,SO, originating from Tambora’s 1815 eruption 


amounted to 150-200 in units of 10° tons, while Krakatau’s 1883 eruption yielded 50-55, in 
the same units. Thus, on these estimates, the mass of H,SO, connected with Tambora was 
something like three to four times greater than that associated with Krakatau. Kondratyev et 
al. (1983, Table 1) do not mention the authors of the foregoing estimates so it is not possible 
to weigh their arguments. A strong point in favour of the estimates of Langway et al. (1988) 
is that their acidity determinations were made at more than one place (South Pole; Créte and 
Dye 3 in Greenland), and that they used data on well-documented, atmospheric nuclear 
explosions by the United States and the Soviet Union. For example, the type of activities, 
amounts released (“source strength’), decay characteristics of radioactive particles, place and 
time of explosion, etc. are known. Moreover, in the documented cases information is available 
on deposition of the total 6-activity. Possible shortcomings of the method used by Langway . 
et al. (1988) may be that in the eruptions of Tambora and Krakatau, the nuclear explosions 
occurred at different times and places: the global transport and dispersal processes may have 
varied from one case to the other. 

. Apparently Trondheim, Norway is exceptional. Birkeland (1949, Table 1, pp. 18-21) shows 
that the lowest annual temperature of the whole series 1761-1946 (3°C) was reached in 1784. 
That low temperature may have been caused by the eruption of the Icelandic volcano Laki in 
June 1783 (see Wood, this volume). Kondratyev et al. (1983, Table 1) provide an estimate 
that, in the aftermath of the eruption, some 10* tons of stratospheric H,SO, was produced. 
Though that mass was 33 to 50% less than the yield of Tambora in 1815, it was the largest 
yearly signal of ice-core acidity in the whole series spanning 553 to 1972 measured at Créte, 
Greenland (Hammer et a/. 1980, Figure 1). No doubt, the size of the signal was due, at least 
in part, to the proximity of Créte and Laki (650 km). Also in the central England record 
(Manley 1974, pp. 393-398), one of the lowest annual temperatures occurred in 1784, and it 
is worth noting that central England is about the same distance from Iceland as Oslo. 
Similarly, the series of annual temperatures 1680-1980 for Paris (Dettwiller 1981, p. 107) 
show that one of the lowest occurred in 1784. However, we cannot be sure that the low 
temperatures of central England and Paris in 1784 were connected with Laki’s eruption. On 
the other hand, the cold winter of 1783-84 at Stockholm (Hamberg 1906, p. 12), where WNW 
to NW winds are frequent, may have been connected with Laki’s eruption. 

. Crop failure caused by drought is infrequent in the Nordic countries. More commonly crop 
failure is due to cold springs or springs and summers; or a superfluity of rain in summer. 

. The Danske Statstidende was the forerunner of the present Copenhagen daily Berlingske 

. A comparison of data of the meteorological stations of the Russian Empire operating in 1812 
(Kiev, Reval (Tallinn), Riga, St. Petersburg (Leningrad), Vilnius, V6yri and Warsaw) 
indicates that, although the year on the whole was rather cold, October was somewhat milder 
than the long-term average. About mid-October Napoleon’s Grande Armée began its retreat 
from Moscow. November was relatively cold and December was particularly cold. Much of 
the heavily decimated army reached Vilnius in December amidst a cold wave, with air 
temperatures around -20°C. The army had crossed the Niemen River in June on its way to 
Moscow (almost on the day of the month when the Germans crossed the same river into the 
Soviet Union in 1941). When the Grande Armée entered Russia, officers and men were 
well-dressed and well-fed; when they reached Vilnius in December on their way back, the 
remainder were in tatters and suffering from exposure and famine. 


7. It is of interest to note that in 1888, long before the volcanic particles could be sampled and 
long before various remote-sensing techniques became available, the Krakatau Committee of 
the Royal Society of London suggested that the volcanic ‘dust’ consisted of ‘condensed 
gaseous products of the eruption (other than water) such as sulphurous and hydrochloric acid’. 
(Deirmendjian 1973, p. 274). Dyer (1971, p. 757) cites a personal communication from 
Mossop to the effect that after Agung’s eruption in 1963, high-flying aircraft were found to 
be thickly coated with a water-soluble material, most likely sulphuric acid. 

8. Hofmann and Rosen (1983a, p. 315) state that volcanic particles they observed over Texas 
after El Chich6n’s eruption (Mexico, spring 1982) were primarily in two layers. At 25 km, 
the aerosol was composed of an 80% H,SOQ, solution, while the lower aerosol layer at 18 km 
was composed of a 60-65% H,SO, solution. 

9. The Danish tenant-farmer and Parish Executive Officer Sgren Pedersen (see the section 
Denmark - Grain Supply) wrote that 1816 was a productive [‘fertile’] year and that high prices 
were obtained for grain "as this commodity was very much coveted in foreign lands". In other 
words, the harvest failure of the countries affected by Tambora’s eruption profited the Baltic 
lands which were not afflicted and which were able to export surplus grain. A passage from 
a report on grain production and export in another country facing the Baltic Sea (Méglinsche 
Annalen der Landwirtschaft, Berlin, 1818, Volume II, pp. 540-552) is worth quoting. The 
report (p. 546) ‘Landwirtschaftliche Bericht aus Meklenburg-Schwerin vom ersten Julius 1818’ 
translates as follows: "After so many dreary years the last two years [1816 and 1817] were 
very favourable for Mecklenburg. Everything cheers up and gets a bright appearance and 
looks forward with confidence to the future. The import ban of England, which was so 
harmful initially to our country, has now turned to our advantage. It drove the grain prices 
high in England, and, consequently, our prices ... The rise in prosperity can also be seen in 
the surplus of capital at low interest rates." 


The Years without a Summer in Switzerland: 1628 and 1816 

Christian Pfister! 

Years without Summer - 1628 and 1816° 

These two summers embody extreme cases of a type that occurred quite frequently within the 
Little Ice Age. In 1816, the thermometer remained well below the average of many years in all 
three summer months: with 13.4°[C] that June just reached today’s average temperature for May, 
and that July and August with 14.9° and 14.8°, respectively, reached the average of a very cold ~ 
June now. Without exception, the other months were extremely cold as well; moreover, April- 
May and July-August were extremely wet: in Bern there was precipitation on 52 days in the 
summer months. Lakes Neuenburg, Murten, and Biel [Neuenburgersee, Murtensee, Bielersee] 
formed a contiguous body of water throughout the summer. In all summer months snow reached 
down to altitudes around 1000 m; on 6 June [1816] the snowcover extended from the Rigi down 
to Weggis. In the Rhine Valley at Biinden [Biindner Rheintal] snow persisted from 8 to 10 June 
down to the 1050 m level; every two weeks it snowed farther down. In the night from 30 to 
31 August Salis heard avalanches thundering down from the Calanda massif. Cattle could not be 
let out onto the higher alpine meadows (Pfister 1976; p. 84 seq.). 

At the earliest locations in the central plain [Mittelland] the rye harvest could not start until early 
August, the cherries in the Rhine Valley at Biinden [Biindner Rheintal] ripened only by the end 
of that month, and the vines had barely finished flowering by that time in some places. At the 
end of September the potato fields were as green as in July. In the higher locations potatoes, 
wheat and oats could not mature any more and ended up covered by snowfalls starting in late 
October. The green grapes had to be picked out of the snow around St. Martin’s Day 
[31 October]. In the Siggental in Aargau they were dragged into a fountain trough and worked 
over with a flail before dumping them into a wine-press. From a comparison of proxy data and 
weather reports (Table 1/19) we can assume that the summer of 1628 can hardly have been any 

'  Historisches Institut, Universitat Bern, Engehaldenstrasse 4, CH-3012 Bern, Switzerland. 

2 Excerpts from Klimageschichte der Schweiz 1525-1860 (Band I:140-141. 1984. Verlag Paul Haupt, Bern and 
Stuttgart) translated and included here with the author’s permission. 


1816 (Deviations from the Mean of Many Years). 

Start of vine flowering 

Driving cattle to alpine 

Beginning of rye harvest 

Beginning of grape harvest 

Must yields (deviation 
from trend) 

Wine quality 

Latewood density 
(index values) 

Snowfall events at levels 
below approximately 
2000 m (June-August) 

11 July (+33 days) 

8-10 July 

31 July (+17 days) 

29 October (after 
frost!) (+22 days) 


very poor 
846 (-154) 
"Covered the mountain, 

i.e., the Engstligenalp 
(1964 m) within seven 

Table 1/19: Comparison of Proxy Data and Weather Reports in the Summers of 1628 and 

8 July (+29 days) 

6 July 

13 August (+32 days) 

9 November 
(+32 days) 

-99 % 

very poor 

809 (-191) 


weeks (i.e., to the 
end of August) 
23 times”! 

Sources (unless quoted otherwise in the Climhist documentation): 

Bartschi, 1916, p. 3. 
2 Strasser, 1890, p. 190. 
3 Observations in the Calanda Region (Pfister, 1976, p. 32). 


Pfister, C. 1976. Die Schwankungen des Unteren Grindelwaldgletschers im Vergleich mit 
historischen Witterungsbeobachtungen und-messungen. Jn: Messerli, B. et al.: Die 
Schwankungen des Unteren Grindelwaldgletschers seit dem Mittelalter. Ein interdisziplinarer 
Beitrag zur Klimageschichte. Zeitschrift fiir Gletscherkunde 11(1):74-90. 


Climatic Conditions of 1815 and 1816 from Tree-Ring Analysis in the 
Tatra Mountains 

Zdzistaw Bednarz' and Janina Trepinska’ 


Dendroclimatological analyses carried out on the stone pine (Pinus cembra L.) in the Tatra 
Mountains indicate that the first half of the nineteenth century was abnormally cold and rainy. 
In this period - the last episode of the Little Ice Age, the year 1816 (quoted by climatologists as ~ 
"the year without a summer") claims special attention. Tree-ring widths of the stone pine are 
abnormally narrow due to low air temperature and high rainfall during the summer of 1816. 
Severe social and economic consequences (e.g., the food deficit and rising prices) of the climatic 
anomalies of 1816 in the Polish Carpathian Mountains were reported by Polish newspapers. The 
climatic anomalies of 1815 and 1816 are probably associated with the world’s greatest recorded 
eruption, that of the volcano Tambora on the island of Sumbawa, Indonesia during April 1815. 
Similar climatic anomalies (cold and rainy weather) were observed in 1912 and 1913 after 
eruption of the volcano Katmai in Alaska in June 1912. Also, in this case, the exceptional 
climatic character of these two years was associated with abnormally narrow rings in European 
high mountain trees and a rapid drop in maximum density of the wood. 


Many dendroclimatological investigations from different parts of the world indicate that tree-ring 
widths can be used as source of information about past climatic changes (Fritts 1976; Eckstein 
and Aniol 1981; Schweingruber 1983). The method is particularly sensitive to extreme climatic 
conditions that exert a decisive effect on tree growth. The alpine timberline is especially 
informative to dendroclimatologists, because that is where the main climatic factor limiting tree 
growth is low air temperature during a short growing season (Eckstein and Aniol 1981; 
Schweingruber 1983; Bednarz 1984). 

Among trees of the upper forest border in the Tatra Mountains, stone pine (Pinus cembra L.) 
deserves special attention. Analysis of the relationship between tree-ring widths of the stone pine 
from a few sites in the Tatras with average monthly air temperature (1911 to present), showed 
a close relationship between tree-ring widths and June-July temperatures. Heavy precipitation in 
those months also limits tree growth (Bednarz 1984). On the basis of these relationships, June 
to July temperatures in the Tatra Mountains were reconstructed for 1741-1911. The reconstructed 
June and July mean air temperatures show a few cold periods (Figure 1). The greatest and longest 
cold period, justly called the last episode of "Little Ice Age", occurred in the Tatras and in the 
Alps during the first half of nineteenth century. The coldness is documented by advances of 
Alpine glaciers then (LeRoy Ladurie 1967; Messerli et al. 1978; Bircher 1982) (Figure 1). 

' Department of Forest Botany and Nature Protection, Agricultural University, 31 425 Krakéw, Al. 29-Listopada 

46, Poland. 


Yhéoo fled 

1740 50 60 70 80 90 1800 10 20 30 40 50 

Figure 1: Reconstructed mean air temperatures for June-July from 1740 to 1850. A- Tatra Mountains, 
B- Alps (summer data from Eckstein and Aniol 1981). Known glacial advances are indicated 
by diagonal hatching. 

Climatic Conditions in 1815 and 1816 in the Tatra Mountains 

Toward the close of the Little Ice Age, 1816 (sometimes called "the year without summer") 
deserves special attention (Stommel and Stommel 1979). Severe social and economic 
consequences of the climatic anomalies of the years 1815 and 1816, in the Polish Carpathian 
Mountains were reported in Polish newspapers. Apart from news about food deficits and rising 
prices, remarks on weather conditions can also be found. For example, Gazeta Lwowska (The 
Lw6éw Newspaper) for 1816 reports: 

"In the information on this year’s harvest received from Biala on October 13, 1816, 
we read that for over 30 years the inhabitants have not remembered such bad 
weather conditions. From autumn 1815 rains did not stop almost to the present day, 
except for a few weeks of better weather from the end of August to the middle of 
September" (Sadowski 1980). 

Also in western Europe (Germany, France and neighbouring countries) the weather was 
abnormally bad. After the long and frosty winter of 1815-16, an exceptionally rainy summer 
began. From May 1816 rains did not stop until August, damaging the crops. -Thus, food was 
imported to western Europe from Russia and North America (Borisenkov and Pasuckij 1988). 
In contrast to the Tatras and Alps, in lower areas of Poland (Cracow, Warsaw) the weather in 
1815 and 1816 was normal or almost normal (Sadowski 1980; Trepinska 1988). In the period of 
instrumental recording which began in the Tatra Mountains in 1911, a similar drastic drop in air 
temperatures is noted for 1912-26. 


Volcanos and Climate 

The years 1912 and 1913 are of particular interest. Owing to a deep growth depression observed 
in all investigated Tatra stone pines, 1912 and 1913 are particularly useful as marker years for 
dendrochronologists (Bednarz 1975). Like 1816, 1912 and 1913 were characterized by 
exceedingly low air temperatures and heavy rainfall in summer. The exceptional character of 
these two years was stressed by Schweingruber (Schweingruber et al. 1979), who observed that 
a rapid drop in the maximum density of wood in trees from high mountain sites throughout 
Europe coincided with the June 1912 eruption of Katmai in Alaska. The influence of this eruption 
on growth processes of mountain trees is probably due to the fact that from the end of June until 
August 1912 the air transparency dropped to 0.56, the normal being 0.77 (Kalitin 1938). Such 
a phenomenon was observed for two years. The decrease of air transparency resulted in a more 
than 20% reduction in direct solar radiation (Griggs 1928; Lamb 1970; Budyko 1971). Many 
climatologists (Lamb 1970; Budyko 1971; Stommel and Stommel 1979; Stothers 1984; Kelly and 
Sear 1984) associate the weather anomalies of the first half of the nineteenth century with a 
period of increased volcanic activity. 

Among severe volcanic eruptions, particular attention should be paid to that of Tambora in April 
1815 - "the largest and deadliest volcanic eruption in recorded history" (Stothers 1984). 


Bednarz, Z. 1975. Geographical range of similarities of annual growth curves of stone pine 
(Pinus cembra L.) in Europe. In: Bioecological Fundamentals of Dendrochronology: 
Symposium Materials. T.T. Bitvinskas (ed.). XII International Botanical Congress, 
Leningrad, July 1975. pp. 75-83. 

. 1984. The comparison of dendroclimatological reconstructions of summer temperatures 
from the Alps and Tatra Mountains from 1741-1965. Dendrochronologia 2:63-72. 

Borisenkov, E.P. and V.M. Pasuckij. 1988. Tysiaczletniaja letopis nieobyczajnych jawlenij 
prirody. Mysl, Moskva. 522 pp. 

Bircher, W. 1982. Zur Gletscher - und Klimageschichte der Saastales; Glazialmorphologische und 
dendroklimatologische Untersuchungen. Physische Geographie 9:1-233. 

Budyko, M.I. 1971. Climate and Life. Leningrad. 470 pp. (In Russian. Also published in English 
by Academic Press, New York, 1974). 

Eckstein, D. and R.W. Aniol. 1981. Dendroclimatological reconstruction of summer temperatures 
for an alpine region. Mitteilungen der forstlichen Bundesversuchsanstalt 142:391-398. 

Fritts, H.C. 1976. Tree Rings and Climate. Academic Press, London. 567 pp. 
Griggs, R.F. 1928. Das Tal der zehntausend Dampfe. Brockhaus, Leipzig. 334 pp. 

Kalitin, N.N. 1938. Aktinometriia. Gidrometeoizdat, Leningrad-Moskva. 324 pp. 


Kelly, P.M. and C.B. Sear. 1984. Climatic impact of explosive volcanic eruptions. Nature 

Lamb, H.H. 1970. Volcanic dust in the atmosphere, with a chronology and assessment of its 
meteorological significance. Philosophical Transactions of the Royal Society of London 

LeRoy Ladurie, E. 1967. Histoire du climat depuis l’an mil. Flammarion, Paris. 381 pp. 

Messerli, B., P. Messerli, C. Pfister and H.T. Zumbuhl. 1978. Fluctuations of climate and 
glaciers in the Bernese Oberland, Switzerland, and their geoecological significance, 1600 
to 1975. Arctic and Alpine Research 10:247-260. 

Sadowski, M. 1980. Czy rzeczywiscie "rok bez lata". Problemy 5(410):33-36. 

Schweingruber, F.H. 1983. Der Jahrring: Standort, Methodik, Zeit und Klima in der 
Dendrochronologia. Verlag Paul Haupt, Bern und Stuttgart. 234 pp. 

Schweingruber, F.H., O.U. Braker and E. Schar. 1979. Dendroclimatic studies on conifers from 
central Europe and Great Britain. Boreas 8:427-452. 

Stommel, H. and E. Stommel. 1979. The year without a summer. Scientific American 

Stothers, R.B. 1984. The great Tambora eruption in 1815 and its aftermath. Science 224:1191- 

Trepinska, J. 1988. Many years run of air pressure and temperature in Cracow against the 

background of their variability in Europe. Uniwersytetu Jagiellonskiego. (Roztrawy 
habilitacyjne; Series nr. 140) Krakéw. pp. 140-169. 


Major Volcanic Eruptions in the Nineteenth and Twentieth Centuries 
and Temperatures in Central Europe 

Vladimir Brizek’ 


This paper deals with temperatures in Prague (1771-1987) as related to the major volcanic 
eruptions (11) in the nineteenth and twentieth centuries, particularly that of Tambora in April 
1815. The circulation, which is the primary cause of temperature changes, was observed with the — 
help of the synoptic periods of the German Hess-Brezowski classification, known as the 
Grosswetterlagen, for 1881 to 1987. Altogether 24 follow-up months after 11. eruptions were 
studied, i.e., 264 months. However, positive temperature deviations (in Klementinum) 
predominate over negative ones 154 to 110. Therefore it cannot be said that volcanic eruptions 
are followed by cooling in central Europe within the next two years. 

Observing the circulation with the help of Grosswetterlagen, we can see in the absolute majority 
of cases, that the drop in temperature in individual months was caused by advection of cooler air 
masses into central Europe. Because of the geographical extent of the Grosswetterlagen 
classification that covers an area larger than Europe, we may conclude that western Europe also 
had cooler weather for the same reasons. 

Therefore I assume that the cooler summer of 1816 in western as well as central Europe 
following the Tambora eruption on Sumbawa in April 1815 was caused by cold air mass 
advection from the north, especially to western Europe where cooling was more marked than in 
our area. The general circulation plays the primary role. 


Atmospheric pollution by volcanic dust that might cause temperature decreases in certain regions 
on our planet is a problem, and the focus of considerable attention. It is assumed that enormous 
pollution might result in fundamental climatic changes, like the ones due to the explosions of 
huge meteorites in the Earth’s geological past. 

This paper deals with temperatures in Prague as related to major volcanic eruptions in the 
nineteenth and twentieth centuries, particularly that of Tambora on Sumbawa in April 1815. The 
circulation, which undoubtedly is the primary cause of temperature changes, was observed with 
the help of the synoptic periods of the German Hess-Brezowski classification, known as the 
Grosswetterlagen, for 1881 to 1987. Furthermore, a series of monthly temperature means 
measured at the Prague Klementinum station between 1771 and 1987 were used. 

Processing Method 
When studying the problem, those eruptions were selected during which at least 0.5 km* of 

material was ejected. In the resulting table (Table 1), the month of the eruption was denoted as 

! Czech Hydrometeorological Institute, Na Sabatce 17,143 06, Prague 4, Komofany. 


0, and for months -1 to +24, deviations from the monthly temperature means for each of the 11 
selected events were recorded. For each month, average deviations were determined for volcanoes 
to the west as well as to the east of central Europe and, finally, their mean. 

Table 1A: Data on Major Volcanic Eruptions in the Nineteenth and Twentieth Centuries. 

Volcano Locality Date of Eruption Material Ejected 

1. Tambora Sumbawa April 1815 (80+) 

2. Coseguina Nicaragua 20 Jan. 1835 25 

3. Mount Shasta California December 1860 2 

4. Krakatau Indonesia 26 Aug. 1883 17 (approx.) 

5. Santa Maria Guatemala November 1902 10 

6. Katmai Alaska 2 June 1912 15 

7. Anizapu Chile summer 1932 20 

8. Bezimyanny Kamchatka 20 March 1956 2 

9. Vesuvius Italy 28 Oct. 1979 0.5 

10. St. Helen’s Washington 18 May 1980 jl) 

11. El Chichén Mexico 28 March 1983 0.6 

For months showing a negative temperature deviation of at least -1°C, I determined the number 
of days on which the synoptic situation occurred that brought cooler weather, or below-average 
temperatures in that month. Of the overall number of 29 Grosswetterlagen synoptic periods we 
formed, e.g., for the summer months, a group of Wc, Ws and Ww situations together with all 
(nine) northern and northwestern ones, including troughs and lows in central Europe. These 
represent a cooler circulation. Similarly, for winter months, we observed the occurrence of 
eastern, northeastern and northern situations accompanied by below-average temperatures, etc. 
The data were tabulated, together with the name of the volcano, date of eruption and quantity of 
ejected material, as well as temperature deviations for individual months (as measured at Prague 
Klementinum), and the average monthly deviations for eruptions to the west and east of central 
Europe and their mean. 


Among the 11 episodes reviewed (except for the eruption of Laki on Iceland in 1873 with 12.5 
km? of ejected material, where the month of the event was not known to the author), the Tambora 
episode was followed by a series of eight months with negative deviations - from May 1816 to 
December 1816. The relatively largest deviations were found for June (-1.2°C), July (-1.2°C), 
August (-1.7°C) and September (-1°C). In the remaining months, the deviations ranged from -0.3 
to -0.9°C. However, in summer months, these negative deviations do not mean anything out of 
the ordinary. The absolute extremes of the Klementinum series are -4.7°C in 1923 for June, 
-4.6°C in 1771 for July, -3.2°C in 1833 for August and -5.0°C in 1912 for September. The 


Table 1B: Monthly Temperature Deviations (°C) at Prague Klementinum as Related to Major 
Volcanic Eruptions (Table 1A). 













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' Months are given before and after (-1 to +24) the eruption of Tambora in April 
1815 (0 in this scheme). 

2 Volcano numbers are keyed to names and other data in Table 1A. 

3 Pw and Pe denote the arithmetic mean of temperature deviations for cases of seven 

volcanoes situated to the west and four volcanoes situated to the east of central 

Europe, respectively. P denotes the arithmetic mean of temperature deviations for 

all 11 volcanoes studied. 


summer of 1816, with its average temperature deviation of -1.37°C at Prague Klementinum, was 
not extremely cold, unlike western Europe. The deviations found for the months under study do 
not nearly approach the above extremes. Regarding the other cases, there does not appear any 
uninterrupted, longer series of negative deviations, except for Vesuvius in 1979, following the 
eruption of which seven months (April to October 1980) exhibited a negative deviation. However, 
only three months showed a deviation from -1.9 to -2.6°C, while the others deviated by only -0.1 
to 0.7°C. So, the figures for 1816 are not extraordinary. 

Altogether 24 follow-up months after 11 eruptions were studied (i.e., 264 months). However, 
positive temperature deviations predominate over the negative ones by 154 to 110. Therefore it 
cannot be said that volcanic eruptions are followed by cooling in central Europe within the next 
two years. 

Nor do the average monthly temperature deviations, or their means, show any cooling down in 
the cases of volcanic eruptions both to the west and east of central Europe; rather the opposite. 
In the western case, 18 deviations were positive and six negative, whereas in the eastern case 17 
were positive and seven negative. In all 11 cases, 17 positive and seven negative deviations were 
recorded within two years after the month of the eruption. 

Let us observe the circulation with the help of the Grosswetterlagen synoptic periods (since 
1881), in the interval of -1 to +24 months surrounding the month in which the eruption took 
place, including that month. In the following, months showing a deviation of at least -1°C or 
more will be considered. In the eight cases under review, negative deviation ranging from -1 to 
-3.4°C occurred five times in the month of eruption. In the preceding month, a negative deviation 
from -1 to -9.8°C also occurred five times. Counting the number of days with cooler synoptic 
periods in a given month, we may conclude that in the eruption months there were 24 to 27 days 
with cooler circulation; the figure is 18 to 28 for the month prior to the eruption. Among the 192 
months which followed the eruptions, in 48 instances the temperature deviation was at least -1°C. 
Of these, only seven months included a mere seven to 12 days with cooler circulation. In the 
remaining 41 cases, the number of such days was 15 to 30. 

In the absolute majority of cases, the drop in temperature in individual months was caused by 
advection of cooler air masses into central Europe. Because of the geographical extent of the 
Grosswetterlagen classification that covers an area larger than Europe, we may conclude also that 
the western European continent had cooler weather for the same reasons. 

Probably the cooler summer of 1816 in western as well as central Europe following the Tambora 
eruption in April 1815 was caused by atmospheric circulation that brought cold air masses from 
the north and northwest, especially to western Europe where cooling was substantially greater 
than in our area. Where climatic change is concerned, atmospheric pollution by volcanic dust 
perhaps plays a secondary role, the general circulation of the atmosphere being of prime 

This work has been inspired by the questions connected with the cold summer of 1816 in western 

Europe following the 1815 Tambora eruption. This topic was the subject of an international 
conference "The Year without a Summer? Climate in 1816" held in Canada in June 1988. 


On the basis of the Klementinum temperature series and the Hesse-Brezowski classification of 
synoptic periods, I conclude that the summer cooling of 1816 in central and western Europe was 
caused primarily by atmospheric circulation - volcanic dust being a factor of secondary 


I acknowledge the effort of Dr. Zdenék Kukal (Central Geological Institute of Prague) for 
collecting and making available the required data on volcanoes. 



Climate over India during the First Quarter of the Nineteenth Century 

G.B. Pant', B. Parthasarathy’ and N.A. Sontakke’ 

A detailed examination of climatic conditions over the Indian subcontinent for 1801-25 has been 
made usi